# Emplacement of a lamprophyric sheet swarm under ductile transpression in continental crust. Insights from a Variscan magmatic complex in the West Armorican belt, Western France

Bernard Le Gall, Martial Caroff

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Bernard Le Gall, Martial Caroff. Emplacement of a lamprophyric sheet swarm under ductile transpression in continental crust. Insights from a Variscan magmatic complex in the West Armorican belt, Western France. Tectonophysics, 2023, 859, pp.229899. hal-05268875

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# Emplacement of a lamprophyric sheet swarm under ductile transpression in continental crust. Insights from a Variscan magmatic complex in the West Armorican belt, Western France

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5 Bernard Le Gall¹*, Martial Caroff¹

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7 ¹Géo-Ocean, Univ. Brest, CNRS, Ifremer, UMR6538, F-29280 Plouzané, France

8 *Corresponding author

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10 **Keywords:** lamprophyre intrusive sheets, ductile strain, host-rock anisotropies, transpressional setting, West Armorican Variscides.

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13 **Abstract**

We present structural data from a Variscan magmatic complex of dominantly lamprophyric (kersantite) and associated felsic (microgranodiorite) sheets - the West Armorican swarm (WAS) - injected into Paleozoic metasedimentary series at the western extremity of a transpressional system in the Armorican belt, Western France. The comparative structural analysis of nearly 100 individual intrusive sheets and their host-rocks demonstrates that the WAS complex is a syntectonic magmatic structure which intruded during the onset of regional shortening via successive pulses alternating with three strain episodes (D₁-D₂-D₃). The moderate/high dipping attitude of a great majority of intrusions, along with their prominent orientation parallel to the regional cleavage (S₁), i.e. orthogonal to the direction of shortening, challenge theoretical predictions about magma-filled fractures in contractional settings. These apparently atypical features are reconciled in a kinematic model which emphasizes the influence of both host-rock planar anisotropies (S₁ cleavage and D₃ shear patterns) and
structural inheritance on magma pathways, and instead hypothesizes the negligible role of the ambient stress field. Comparisons are attempted with syncollisional lamprophyre systems present in the European Variscides, and more especially in the SE Bohemian massif.

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# 1. Introduction

Development of collision-related orogenic belts is generally accompanied by magmatic processes which express in various ways as a function of many parameters including: (i) the magma petrology and rheology, (ii) the mode of magma emplacement (intrusive vs extrusive and plutons vs sill-dyke swarms), (iii) the depth of magma emplacement (infra- vs supra-crustal levels) and (iv) the strain/magmatism timing relationships (syn-, late-, post-tectonic/collisional systems). The combination of these parameters leads to a great variety of magmatic rock-types, the most of which have been widely documented worldwide. However, syncollisional (sensu lato) sill/dyke swarms injected during shortening through mid/upper-crustal levels have received so far very little attention. Those published in literature chiefly concern: (i) intrusions injected in lower (Cagnard et al., 2004) or shallow crustal levels (Foster et al., 2001; Galland et al., 2003; 2007a, b; Mazzerini et al., 2010; Menand et al., 2010; Montanari et al., 2010a, b; Ferré et al., 2012; Tibaldi et al., 2017), and (ii) late-post collisional intrusions emplaced during extensional/transtensional stages that commonly accompany the gravity collapse of orogenic mountains (Gapais et al., 1993; Zeitlhofer et al., 2016).

The limited number of investigated synorogenic sill/dyke networks likely reflects the generally agreed assertion that opening of vertical crack/fracture networks as potential magma feeder conduits is made difficult in contractional tectonic regimes because of the horizontal attitude of the maximum principal stress (Anderson, 1951; Glazner, 1991; Hamilton, 1995; Watanabe et al., 1999). Consequently, the mechanisms of injection and propagation of dyke intrusions in compressional environments, as well as the way the strained host-rocks accommodate their
51 emplacement, are not yet fully understood (e.g., Schoffield et al., 2012; Bertelsen et al., 2018; 52 Kjoll et al., 2019).

53 The aim of our work is to partly fill this gap by presenting structural results about a magmatic 54 swarm emplaced during the Variscan orogeny in a transpressional setting at the western 55 extremity of the Armorican belt, Western France (Fig. 1d). The so-called West Armorican 56 swarm (WAS) is a lamprophyric complex composed of kersantite and microgranodiorite sheet- 57 like intrusions (Caroff et al., 2021) cutting through deformed Paleozoic metasedimentary series 58 in the Brest syncline (Figs. 1d, 2). The main goals of our field-based study are: (i) to characterize 59 the main geometric attributes (strike, dip, width, spacing, ..) of ca. 100 individual intrusions 60 continuously exposed along coastal sections in the Brest bay (Fig. 2), (ii) to precise their spatial 61 and temporal relationships with the polystrained metasedimentary host-rocks, and (iii) to 62 qualitatively discuss still challenging issues about the competing role played by various 63 physical parameters on the trajectory and final shape of individual intrusions. Lastly, the 64 structural significance of the WAS complex is discussed with regards to the lamprophyre and/or 65 syncollisional intrusive sheet networks present: (i) in the Armorican belt and (ii) in the entire 66 European Variscides, with emphasis on the Bohemian systems which display a wide range of 67 temporal patterns and structural settings. 68

69 **1. Geodynamical and geological settings**

70 The Variscan belt in Europe is composed of a mosaic of continental blocks which collided 71 separately, after multiple and diachronous oceanic subductions (e.g. von Raumer et al., 2014). 72 That resulted in an intricate pattern of orogenic segments displaying their own tectonic and 73 magmatic histories (e.g. Kroner and Romer, 2013). During Carboniferous times, the oblique 74 convergence of the Gondwana and Laurussia plates induced a large-scale transpressional setting 75 in the intervening Armorica micro-plate, at the western extremity of the orogeny (Gumiaux et
al., 2004; Ballèvre et al., 2009). Transpression chiefly expressed by dextral ductile wrenching along the North and South Armorican shear zones (NASZ, SASZ) and their satellite structures (Elorn, Monts d'Arrée and Montagnes Noires shear zones, ESZ, MASZ and MNSZ, respectively, in Fig. 1d), in response to a NW-SE-directed shortening (Gapais and Le Corre, 1980; Darboux and Le Gall, 1988; Le Gall et al., 1992; Gumiaux et al., 2004; van Noorden et al., 2007). These crustal-scale discontinuities currently separate high-grade metamorphic and plutonic terrains in the North and South Armorican domains (in addition to Cadomian basement rocks in the North) from a thick pile (>5 km) of Late Proterozoic-Paleozoic low-grade metasedimentary rocks, pierced by granitic intrusions, in a central domain (CAD, Fig. 1d). At the western extremity of the CAD (WCAD), the moderate amount of strain resulted in a map-scale syncline pattern (Brest and Châteaulin) which was the locus of the WAS magmatic complex under study (Figs. 1d, 2). The age of the onset of strain in the WCAD is approximated to ca. 315 Ma by correlations with the age yielded by synkinematic granites (Saint-Renan, Huelgoat) along the NASZ and MASZ, respectively (Fig. 1d) (Le Gall et al., 2014; Ballouard, 2016). These tectono-magmatic events are nearly synchronous to the peak of metamorphism in the adjoining LMD (Figs. 1d, 2) (Schulz, 2013; Marcoux et al., 2009; Authemayou et al., 2019). The Brest syncline extends as a 25 x 50 km² bowl-like fold, south of the Elorn bounding shear zone (Fig. 2). Its progressive narrowing westwards (from 50 to 30 km) is accompanied by the counterclockwise rotation of the regional structural grain from N90°E to N60°E. The ca. 3 km-thick succession of Paleozoic metasedimentary series involved in the Brest syncline ranges in age from Ediacaran/Cambrian (Brioverian) up to Late Devonian (Fig. 2b) (Chauris et al., 1980; Babin et al., 1982). Their lithological content can be simplified into two contrasted categories. A ca. 1200 m-thick lower lithostratigraphic part (Ordovician-Lower Devonian) is dominated by competent series, including the Armorican Sandstone Fm. (<200 m), the Plougastel Schist and Quartzite Fm. (400 m), the Landevennec Sandstone Fm. (100 m) and the Armorique
101 Limestone Fm. (100 m). They are extensively exposed to the north in a 5 km-wide belt of
102 regional-scale folds whereas only part of them form a monoclinal structure along the southern
103 limb of the syncline (Figs. 2a, c). The axial part of the syncline is occupied by a ca. 1000 m-
104 thick succession of dominantly fossiliferous schist series, Middle-Late Devonian in age, which
105 are the preferential locus of WAS intrusions (Fig. 2a, Chauris et al., 1980; Babin et al., 1982;
106 Castaing et al., 1987; Plusquellec et al., 1999; Darboux et al., 2010). A thorough petrographic
107 analysis of the kersantite and microgranodiorite populations, as well as their geochemical
108 affinities and the nature of the source material, have been recently performed by Caroff et al.
109 (2021). The texture and mineral assemblages of these rocks show limited variations. Kersantites
110 are a fine-grained to moderately coarse-grained material predominantly made up of plagioclase
111 and dark micas, and subsidiary carbonates and secondary quartz. The microgranodiorite
112 material is an aphanitic or quartz-phyric fine-grained leucocratic rock made up of oligoclase,
113 generally albitized or damouritized, quartz, K-feldspar, chloritized biotite and muscovite.
114 Kersantite/microgranodiorite cross-cutting relationships are exhibited at three sites (Rosmellec,
115 Ile Longue, Porsguen; Fig. 2a, Supplementary Material Fig. S1). Though the kersantite material
116 systematically intrudes (and post-dates) microgranodiorite, the cogenetic origin of the two types
117 of magma is argued from petro-geochemical evidence discussed in Caroff et al. (2021).

118 The WAS complex has not yet benefitted from radiometric dating because of unfavorable
119 mineralogy. However, a maximum age is provided by the few kersantite sheets cutting through
120 Visean/early Serpukhovian series in the Châteaulin syncline to the SE (Fig. 2a, Doubinger and
121 Pelhate, 1976). These relative timing data invalidate previous interpretations of the WAS
122 intrusions as syncompressional magmatic structures related to an early (Late Devonian)
123 orogenic event (Rolet and Thonon, 1978). These structural and temporal issues are fully
124 discussed in the present work.
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Addressing the development of the WAS magmatism in the West Armorican Variscan context first requires to discuss the still debatable nature of the structural relationships between the LMD and the WCAD. This issue is illustrated by the sketch structural cross-section in Fig. 1e, elaborated by combining available metamorphic data (Jones, 1994; Schulz, 2013) with recent structural onshore/offshore results (Le Gall et al., 2014; Authemayou et al., 2019).

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North of the pyrophyllite-bearing facies in the Brioverian series, the LMD metamorphic terrains are involved in a southerly-inclined monoclinal pattern on the southern flank of the LMD dome.

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Their mineralogical assemblages typically reflect epizonal-mesozonal P/T conditions that increase gradually northwards from chlorite-bearing facies to biotite, staurolite and sillimanite/migmatite facies in the structurally lower part of the pile. No gap, neither reverse position of metamorphic rock-types is observed in the pile (Jones, 1994; Schulz, 2013). To the south, the Brioverian series pass, without any stratigraphic gap (despite the presence of the ESZ), into Paleozoic pyrophyllite-bearing series in the Brest syncline (Paradis et al., 1983).

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A quite similar tectono-metamorphic arrangement occurs on the eastern edge of the Leon dome (Fig. 1d, Le Gall et al., 2021). The only regional-scale structures so far identified at the LMD-WCAD transition zone are discrete dextral shear zones cutting through either the Brest orthogneiss (Pierres Noires shear zone, PSZ) or Ordovician quartzites of the Armorican sandstone Fm. along the northern flank of the Brest syncline (ESZ; Fig. 1e). In contradiction with previous allochthonous-type models (Balé and Brun, 1986; Rolet et al., 1986; Faure et al., 2008, 2010; Ballèvre et al., 2009), the southern part of the LMD and the adjoining WCAD are here regarded as sub-autochthonous infra- and supra-crustal domains, respectively, forming a quasi-continuous and coherent crustal section on the southern limb of the Leon dome. Their boundary is thus necessarily arbitrary and located here in the Brioverian series (Fig. 1e).

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By further assuming that the lack of Lower Paleozoic sedimentary series in the LMD is not due to erosion (excepted in the South Ouessant island, see Caroff et al., 2016, 2020, for details), it

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is also inferred that in pre-Variscan times the proto-LMD formed an elevated domain in the
uplifted footwall block of the proto-Elorn extensional fault bounding to the north the long-lived
(> 100 Ma) and subsiding WCAD basin.

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## 2. Methodology

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Our work is built on a field-based structural study devoted to both WAS intrusions and their strained metasedimentary host-rocks in the Brest syncline. Because of very poor inland exposures, field investigations have been restricted to the indented coastal section continuously exposed on the NE edge of the Brest bay (Fig. 2c). Our structural data are synthetized on a ca. 20 km-long cross-section perpendicular to the regional structural grain. The spatial and temporal relationships of strain and intrusion development are described in details by a careful field structural analysis. Each type of ductile deformation is defined in terms of geometry, relative chronology and spatial distribution along the reference section. In parallel, the orientation, dip attitude and width of ca. 100 individual magmatic (kersantite and microgranodiorite) intrusions have been measured. Each intrusive sheet is assigned the letter K (kersantite) or µGD (microgranodiorite) and a unique number (1-101 from W to E on the coastal cross-section in the Supplementary Material Fig. S1). Many of them are shown on the figures with their corresponding number. All the sites mentioned in the text are located on the Google Earth map in the Supplementary Material Fig. S1. Our 2D-cross-sectional work does not permit to investigate the length dimensions and map-scale geometry of the intrusions, excepted about the Ile Longue laccolith at the western edge of the syncline, and on a few rocky platforms intermittently exposed at high magnitude low tides (Château, Le Roz, Yelen, Fig. 2a). From cross-cutting relationships with the poly-strained host-rocks, different generations of intrusions are discriminated as pre-, syn-, and post-strain bodies with regards to three successive strain events detailed below. Both the kersantite (88%) and microgranodiorite (12%) material are considered as two homogeneous petro-geochemical populations (Caroff et al., 2021). However,

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at a few localities (North Château), geochemically distinct kersantite intrusions are distinguished because, as being further involved in structures with contrasted styles, they supply supportive evidence for successive pulses of magma at various stages of ductile strain.

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Some clarifications are needed about the terms “sills and dykes” commonly used about sheet intrusions. In most of published papers, the host-rocks consist of unstrained and flat-lying (layered) sediments or volcanics formed in extensional settings. Therefore, because of their conventional definition as intrusions either concordant with or secant to host-rock anisotropies, sills and dykes are routinely considered as horizontal or inclined/vertical magmatic features, respectively. These notions are inadequate about the WAS intrusions under study because, as intruding through folded strata, a given intrusion may lie as a sill-like body in one place and then extends laterally as a dyke-like body. These notions are substituted by the terms “shallowly-inclined (SIS), moderately-inclined (MIS) and highly-inclined (HIS) sheets”, for dipping attitudes (α) arbitrarily set at α ≤ 30°, 30° < α ≤ 60° and α > 60°, respectively. This nomenclature is justified by the fact that most of WAS intrusions, though being interpreted as syn-collisional magmatic structures, chiefly escaped deformation and thus intruded with their present attitudes.

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### 3. Results

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The nature of parameters that might have influenced the emplacement, propagation and final architecture of the WAS magmatic intrusions is addressed below in relation with: (i) the kersantite vs microgranodiorite composition of the magma and (ii) the mechanical behavior of the competent (grits, quartzites, limestones) vs weak (schists) host-rocks during deformation in the Northern vs Southern units. Others potential controlling factors such as tectonic stress and magma properties are not addressed in the present work.

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**4.1. The host-rock ductile strain pattern**

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The 2D-structure of the Brest syncline on the cross-section in Fig. 2c is derived from published geological maps further combined to our own structural field data. It corresponds to a highly asymmetrical and two-fold structure comprising of: (i) a Northern unit, chiefly composed of competent and folded Ordovician-Lower Devonian series, juxtaposed to the south along the SE-directed Rozegat thrust to (ii) highly-sheared Middle-Upper Devonian schist series (Southern unit) that in turn pass southwards into the unfaulted and monoclinal flank of the syncline.

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The overall structure of the Northern unit is dominated by a succession of km-wavelength upright or overturned syncline/anticline folds (D₁ event) (Fig. 2c). Anticline hinge zones are commonly cored by massive quartzitic strata of the Plougastel and Landevennec Fms. Axes of D₁ parasitic folds exposed in the field plunge shallowly (<10° in average), mostly in the SW quadrant (Fig. 3a). D₁ folds are associated with a steep penetrative axial-plane cleavage (S₁), oriented in the range 20-60° (Figs. 3a, b). It is noteworthy that the S₁ planes do not bear any apparent stretching lineations. The Middle/Upper Devonian schist series that locally extend in the associate synclines show a more intricate geometry resulting from overprinting of S₁ planes by D₂ folds, in turn disrupted by a shallowly-inclined network of shear surfaces (D₃ event). A decoupling level, not yet argued in the field, is suspected to occur between the lower (competent) and upper (schist) series (Fig. 2c). Both the D₁ and D₂ fold patterns are disrupted by a system of shear and ramp-thrust structures displaying prominent transport directions to the east and assigned to a D₃ event. A few of these dominantly shallowly-inclined structures are sites of magma injection (Sections 4.2.5, 4.3). From the cut-off geometry of footwall/hangingwall stratigraphic limits, thrust displacements are assumed not to exceed 100 m's, excepted about the Rozegat frontal thrust where they are <500 m (Fig. 2c). The
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northernmost thrust structures might merge at depth along the Elorn bounding dextral shear
zone in a regional-scale flower pattern.

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The Middle-Upper Devonian series in the Southern unit show a more intricate structural pattern
resulting from the superimposed effects of three ductile strain events (Babin et al., 1975). A
shallowly/moderately-dipping prominent cleavage (S₁) represents the axial plane of overturned
to recumbent folds (Fig. 3c), correlated to D₁ structures. Again, the S₁ surfaces do not bear any
stretching lineation. The wide distribution of S₁ strikes, with a peak in the range N50-100°E
(Figs. 3d1, e), is due to a second generation of dominantly upright and shallowly-plunging
(<30° in average) folds, assigned to a D₂ event. The density of S₁ poles in Fig. 3d2 indicates a
mean D₂ fold axis plunging at ca. 30° to N240°E, in agreement with field observations. The D₁-
D₂ multistage fold pattern is disrupted by a network of SE-directed flat-lying shear structures
(event D₃), frequently accompanied by tectonic schist breccias. Their thrust origin is locally
argued, as for example at the Kerloziou site where Upper Devonian schists are tectonically
overlain by Middle Devonian series along a nearly horizontal contact (Fig. 3f; Babin et al.,
1975). Lastly, these early ductile planar fabrics are involved into a succession of a few km's-
wavelength undulations, possibly formed on top of a blind ramp-thrust system involving the
Lower Paleozoic competent series at depth (Fig. 2a).

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The lack of pre-Devonian series currently exposed in the Southern unit may be likely related to
the decoupling level envisaged in the Northern unit (Fig. 2a). Dextral strike-slip tectonics is
restricted to the axial part of the Brest syncline, in relation with the so-called Le Roz shear zone
(Figs. 2a, c). This structural feature further coincides with a major magmatic structure discussed
in Section 4.2.6.

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## 4.2. The WAS intrusive complex

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### 4.2.1. Spatial distribution
Because of limited exposures inland, very little is known about the spatial distribution and map-geometry of WAS intrusions. They are thus drawn as discrete point-like features on published 1: 50 000 geological maps (Fig. 2a; Chauris et al., 1980; Babin et al., 1982; Castaing et al., 1987; Plusquellec et al., 1999; Darboux et al., 2010). New insights are here provided about their amount, location, spacing and density from the 20 km-long coastal reference cross-section (Figs. 2a, 4a), thought to be representative of the WAS complex as a whole.

Along the analyzed section, kersantite are more widespread than microgranodiorite intrusions. The latter are mostly restricted to two main areas (Caro and Le Roz/Ile Longue) which further coincide with the greatest concentration of intrusions in each structural unit. The main cluster of intrusions (x20) occurs in the Le Roz area, i.e. in the axial part of the Brest syncline, from which the intrusion frequency decreases gradually in a symmetric way. In contrast, in the Northern unit, the frequency of intrusions decreases asymmetrically southwards from a maximum peak (x10) in the Caro area. The spatial association of the Le Roz/Ile Longue highly intruded zone with a major dextral shear zone is discussed in Section 4.2.6.

### 4.2.2. Strike and dip

The orientation (dip and strike) of planar intrusions is a key-parameter for identifying realistic geometries with respect to the ambient tectonic stress/strain fields. The strike pattern of nearly 100 individual magmatic sheet intrusions measured in the field is plotted on the rose diagrams in Figs. 4b-d. Intrusion strikes display a dominant regional trend in the range N20-100°E (ca. 65%), with three peaks at N40-50°E, N70-80°E and N90-100°E (Fig. 4b). They rotate clockwise from N30-50°E in the Northern unit (Fig. 4c) to N80-100°E in the Southern unit (Fig. 4d). This intrusion strike changing is fairly consistent with variations of S₁ cleavage orientations in the Northern and Southern units. Subordinate intrusion networks (15%) strike N120°E (in both units) and N160°E (Southern unit) (Figs. 4b-d).
On the diagram in Fig. 4e, the dip pattern (∠) of 95 intrusions is scattered between 10° and 90°, with two peaks in the intervals 30-50° (33%) and 80-90° (19%) for both kersantite and microgranodiorite material, and irrespective of the host-rock lithology (competent vs schist). With regards to the nomenclature adopted above, the HIS (∠ > 60°) and MIS (30° < ∠ ≤ 60°) populations (40% and 50%, respectively) largely exceed those of SIS (∠ < 30°, 10%).

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### 4.2.3. Width

Most of WAS intrusions are tabular bodies showing sharp contacts with the deformed metasedimentary host-rocks. Width measurements (x101), made normal to their walls, have been analyzed by frequency distribution as a function of: (i) their kersantite vs microgranodiorite nature and (ii) their dipping attitude. The measured widths are mostly in the range 0.2-30 m with 65% <5 m (Fig. 4f). The proportion of microgranodiorite vs kersantite is higher for widths >25 m. In contrast, no direct correlations exist between the width values and: (i) the dipping attitude of the intrusions and (ii) the host-rock lithologies. Except for the Le Roz/Ile Longue highly-intruded zones (Section 4.2.6), it appears that the areas of greatest widths (Goasquellou, Rosmellec, Squiffiec; Fig. 2a) do not strictly coincide with areas of high frequency. Neither significant (decreasing) width changes have been observed at increasing distances from the contact of these two zones.

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### 4.2.4. Ring-shaped intrusions

Three kersantite intrusive sheets, cutting through Upper Devonian schists (Porsguen Fm.) on the southern edge of the Château beach, form map-scale elliptical structures (40 x 60 m), aligned over a distance of 300 m along a N140°E axis parallel to their marked elongation (Fig. 5a). The two southern ones (K55, 56) form ring-shaped structures which evoke those more routinely observed in younger volcanic belts (Takada, 1988; Walker, 1999; Pasquare and
Tibaldi, 2007). Each ring-shaped structure is formed by a 6 to 3 m-wide peripheral intrusion, dipping moderately inward on its NE flank and passing into a vertical position on its SW flank. These ring-shaped intrusions are unstrained and show a still preserved and regularly-spaced cooling joint pattern orthogonal to their margins. The third intrusion ($K_{54}$) forms an elongated plug oriented at NE-SW. All the intrusions are wrapped by the regional $S_1$ cleavage.

### 4.2.5. Transition patterns from highly- to shallowly-inclined sheets

Spatial variations of intrusion trajectories from HIS to SIS are observed in two localities. In the South Château area, a kersantite dyke ($K_{59}$), >10 m-wide, passes laterally westwards and upwards into a subhorizontal intrusive sheet, in turn underlain by a top-to-N155° shear surface ($D_3$) and associated asymmetrical fold structures in the underlying Famennian schist host-rocks (Fig. 5b). Along the Caro section (Fig. 5c), a top-to the west and shallowly-dipping shear surface ($D_3$) is cut at high angle by a 4 m-wide microgranodiorite HIS ($\mu GD_{17}$) which displays a 20 cm-thick apophysis along the trace of the shear surface over a distance of 20 m.

### 4.2.6. The Le Roz and Ile Longue magmatic zones

The morphology of the Le Roz peninsula provides a 3D-view of magmatic and tectonic structures along two orthogonal cliff sections and over the 250 x 100 m² adjoining rocky platform (Fig. 6a). On the EW-trending section (Fig. 6b), two distinct tectono-magmatic packages are tectonically juxtaposed. The upper one is composed of a subhorizontal and nearly undeformed microgranodiorite SIS ($\mu GD_{60}$), >20 m in width (its upper margin is not observed). Its lowermost part is disrupted by a series of top-to-the-W shear planes, probably related to a basal thrust contact ($D_3$) with an underlying assemblage of highly-strained Late Devonian schists (Porsguen Fm.) containing closely-spaced kersantite intrusive sheets. >30 % of the rock mass on the 200 m-long section in Fig. 6c is made of 3 to 10 m-wide kersantite sheets (x18,
K<sub>62-79</sub>). It further coincides (in its central part) with a high strain shear zone spectacularly expressed in both the schist host-rocks and the intrusive network. The Porsguen schists occur as tectonic breccias due to intense shearing of the folded (D<sub>2</sub>) S<sub>1</sub> cleavage. The kersantite sheets form a series of disrupted and relatively steep magmatic bodies, with variable dimensions and shapes, wrapped by the brecciated schist host-rocks (Figs. 6a, c). It is noteworthy that the greatest dismembered sheet intrusions are devoid of any internal deformation and still display regular 3D-shapes bounded by the initial magmatic margins (long dimensions) and cooling joint patterns at their extremities (Figs. 6c, d). Smaller kersantite bodies form either blocks, ranging in diameter from a few dm's to a few m's in schist breccias (Fig. 6e), or elliptical boudin-like lenses surrounded by the schist breccia. The great concentration of intrusions along the Le Roz section, along with their dominant steeply-dipping attitude (inferred to be a primary feature), are considered as supportive evidence for a potential magmatic feeder zone (Le Roz) in the axial part of the Brest syncline.

A nearly similar intricate tectono-magmatic pattern is exposed in the Ile Longue area to the SW (Fig. 7). There, two contrasted tectono-magmatic packages, involving only microgranodiorite material, are tectonically juxtaposed. The upper package comprises a >30 m-wide microgranodiorite SIS resting horizontally along a basal sheared contact over flat-lying Upper Devonian schists (Porsguen Fm.) (Figs. 7a, b) (Thonon and Rolet, 1982). This laccolith-like sheet, which represents the only map-scale intrusion in the WAS complex, is cut by a few kersantite HIS. This upper package is thrusted over similar schist series which conversely display a steeply-dipping S<sub>1</sub> cleavage locally involved in a 100 m-wide and N70°-oriented strike-slip shear zone (Fig. 7c). The dextral sense of shearing is deduced from the asymmetry of D<sub>2</sub> folds superimposed on the S<sub>1</sub> cleavage. The shear zone encloses numerous discrete bodies of microgranodiorite material showing various dimensions and shapes, but generally forming elongated lenses surrounded by the S<sub>1</sub> cleavage (Fig. 7b). From their structural similarities, the
350 Ile Longue and Le Roz tectono-magmatic patterns are laterally correlated, being dextrally-offset (>10 km) along a transverse fault (FT) still expressed in the bathymetry of the Brest bay seafloor (Fig. 2a).

353

354

**4.3. Strain/intrusions spatio-temporal relationships**

355 Timing and conditions of emplacement of the WAS magmatic intrusions are constrained by intersection relationships with host-rock planar anisotropies. Emphasis is put below on the strata pattern (S₀), principally about the competent series in the Northern unit, and on the S₁ cleavage (D₁) and flat-lying shear plane (D₃) patterns in both units (Fig. 8a; Supplementary Material Table S1). The moderate amount of intrusive sheets (10 out of 26) occurring parallel with strata in the competent series of the Northern unit indicates that strata boundaries (S₀) play a modest role on sheet trajectory. A quite different structural configuration is observed in the dominantly schist series where about 50% of intrusions (44 out of 100) are seen to follow the trace of the S₁ fabrics, whatever their shallowly- (x4), moderately- (x21) or highly- (x19) dipping attitude. Lastly, a total of 13 intrusive sheets occur along flat-lying shear zones (D₃). Three of them have been described above about the South Château and Caro HIS/SIS transition patterns (Figs. 5b, c), as well as in the Le Roz/Ile Longue shear zone (Figs. 6, 7). One additional example is described below along the Larmor section.

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The relative chronology of magma emplacement vs deformation events has been established about 82 (out of 101) individual intrusions which either pre-date or post-date one of the three strain episodes (D₁, D₂, D₃, Fig. 8b). The chronology of the remaining intrusions (x19) is not constrained. It is noteworthy that intrusive sheets post-dating the late shearing event (D₃) have not been documented in the WAS complex.

373 Only 35 intrusions display a maximum and a minimum time-limits and are thus fully calibrated.

374 Two distinct strain/intrusion temporal patterns are distinguished. (i) Intrusions emplaced at an
early stage of the polyphase strain history (x20) are reported in two distinct, but closely spaced sites.

At the South Château site, the trajectory of the steeply-dipping S₁ cleavage of the host-rocks is moulded around kersantite sheets (x5), whatever their shape as either one elliptical plug (K₅₄), ring-sheets (K₅₅-₅₆) or submeridian sheets (K₅₇-₅₈) (Fig. 5a). The boudinaged structure of the latter, as well as triple junction trajectories of the S₁ cleavage at the extremities of individual sheets, resemble patterns studied elsewhere (Pons et al., 1995) and typically characterize synchronous magmatic and ductile strain (D₁) events. Further south in the Le Roz area, highly-strained (boudinaged) kersantite sheets (x18, K₆₂-₇₉, Fig. 6) are similarly wrapped by the S₁ cleavage which, here, further recorded intense dextral shearing (D₂). These structural features again show that magma injection operated at a relatively early stage of strain, in the time-span D₁/D₂, prior to the thrust emplacement (D₃) of a nearly unstrained microgranodiorite sheet (µGD₆₀, see above Fig. 6). It is thus demonstrated that multistage injections of compositionally distinct magmas occurred at successive stages of the polyphase ductile strain history in the Le Roz area. A similar timing pattern, but involving only kersantite material, is documented at the Château site. The map and cross-section in Fig. 9 show a 30 m-wide HIS (K₅₁), parallel to the vertical S₁ cleavage in Late Devonian schists (Traonliors Fm.) and enclosing schist enclaves (inferred to correspond to S₁), which is thus interpreted as a post-D₁ intrusion. Further west, flat-lying schists occurring in the thrusted (D₃) western limb of a D₂ fold are intruded by a geochemically distinct SIS kersantite (K₅₃) (Caroff et al., 2021), again parallel with the S₁ cleavage, but here in a flat-lying attitude. The post-D₁ emplacement of the two distinct kersantite sheets likely took place prior (K₅₁) and later/during (K₅₃) the folding (D₂) and shearing (D₃) episodes, at a time when the two corresponding magmas were not connected.
Additional supportive evidence for magma injection in the time range (D₁-D₃) are provided by much less intricate structural patterns involving only one individual intrusive sheet. At the
Squiffiec site (Fig. 10b), a microgranodiorite HIS (μGD₃₃) enclosing schist (S₁) enclaves post-dated the D₁ event and is in turn involved into, and thus post-dated by, a D₃ shear thrust structure. A similar sequential timing pattern is observed at the Kerascoet site where a >20 m-wide kersantite SIS (K₈₇) containing a S₁ schist enclave is disrupted along its basal contact by a flat-lying shear thrust (D₃) directed to the north (Figs. 11a, b).

(ii) The latest intrusion network (x7) identified in the WAS complex is genetically related to D₃ shear structures. Two examples have been described above about the South Château and Caro HIS/SIS transition patterns (Figs. 5b, c). In the latter site, folding of the microgranodiorite layering in the SIS apophysis (μGD17 in Fig. 11c) attests that the felsic magma was not yet frozen during the last increment of shearing. But, the most demonstrative example of syn-D₃ intrusion is exposed at the Larmor site. It corresponds to a 3 m-wide kersantite SIS (K₂₃ in Fig. 10c) intruded along a shallowly-inclined and SE-directed thrust fault (D₃) cutting through vertical folded (D₁) quartzite strata (Landevennec Fm.). The kersantite SIS is locally disrupted by satellite D₃ shear zones probably formed at a late stage of the shearing event (Fig. 10d).

The partly-bracketed intrusion population (x44) also supplies fruitful insights as demonstrating that >50% of sheets (x26) post-dated the main ductile events D₁/D₂ and thus intruded at a relatively late stage of the polyphase strain history. This assertion is based on: (i) the presence of schist xenoliths (inferred to contain the S₁ dominant cleavage) within the magmatic bodies, (ii) the slight angular obliquity of the intrusion and the host-rock cleavage (S₁ or S₂), and (iii) the disruption of syncleavage fold structures (D₁ or D₂) by intrusions. The example exposed at the Yelen site (Fig. 10a) shows a 2 m-thick kersantite sheet (K₈₃) cutting at high angle through a D₂ upright fold hinge and then following the S₁ cleavage on the limb of the fold. A different fold/sheet intersection pattern is observed at the Kerdreolet site (Fig. 2a) where the trajectory of a kersantite sheet through folded schist/limestone alternations (Bolast Fm. in Fig. 2b) is offset into two linear segments in the hinge zone of a D₂ upright fold. The upper time-limit of these
post-D₁/D₂ intrusions, and more specifically their temporal relationships with the D₃ strain event, is not constrained (Fig. 8b). In a symmetric way, 21 pre-D₃ intrusive sheets do not display lower time-limits, so that their emplacement prior to the early D₁ strain event cannot be excluded.

429

430 **4.4. Strained vs unstrained intrusive sheets**

Most of WAS intrusions are tabular sheets showing sharp contacts with the strained host- rocks. One of their fundamental structural features is that they escaped pervasive ductile strain (with the exception of the folded µGD₁₇ SIS intruded along a D₃ shear plane). On one other hand, the involvement (or not) of intrusive sheets in brittle strain is a discriminant parameter for identifying two distinct populations. A few strained intrusive sheets have been described above in relation with low-angle shear thrust structures (D₃), but, the great majority of them are restricted to the Le Roz-Ile Longue dextral shear zone (in addition to a few sheets in the South Château area). As shown above in Fig. 6, the spectacularly dismembered (boudin-like) geometry of the kersantite sheet network in the Le Roz shear zone (K₆₃-77) resulted from the combined effect of shear-related stretching, acting in two orthogonal directions. The horizontal stretching is assigned to dextral strike-slip shearing kinematically consistent with the asymmetry of D₂ isoclinal folds. The lack of any penetrative strain in the magmatic bodies, their disruption along cooling joints, and their preserved tabular shape (Fig. 6) indicate that the kersantite sheets have been disrupted during high strain brittle shearing during the D₂ event.

Outside the Le Roz shear zone, the sharp and wedge-shaped geometry of a number of intrusion tips, such as the MIS K₅₆ in the South Château ring-sheet network (Figs. 5a,12b), shows that they also behaved in a brittle way during shortening. Similarly, the brittle mechanical behavior of the strained metasedimentary host-rock during magma injection is locally documented by crack/fracture networks thought to have opened by shear failure ahead of intrusion tips (K₂₄ in
Fig. 12a). Convincing evidence for the brittle behavior of the host-rocks during shortening have been reported above about the late low-angle shear plane pattern (D₃). Additional support is provided by a widely distributed system of narrow (30-40 cm-wide) shear zones following the walls of many intrusions (Fig. 12c). Their magmatic vs tectonic origin can be questioned since magma injection is also known to induce compressional strain in host-rocks (Tibaldi et al., 2017). In the present case, the unique sense of shearing (deduced from the sigmoid geometry of the S₁cleavage) on both sides of the intrusions confirms its tectonic origin. A demonstrative example is shown in Fig. 12c about a kersantite SIS (K₅) at the Ilien ar Gwen site. Senses of shearing are not consistent regionally and might result from local stress disturbances along individual intrusive bodies. This discrete shear zone pattern was likely promoted by high rheological contrasts in the rock volume between the resistant intrusive sheets and the weaker schist host-rocks.

462

463

**5. Discussion.**

The structural results obtained about the WAS intrusions are discussed in two complementary ways. The first step is to evaluate whether they provide valuable constraints to better understanding the way intrusive sheet networks can intrude and propagate through a mid/upper crust recording shortening. The second step is to address the structural significance of the WAS complex in the Variscan orogeny by comparing its mode of injection with: (i) contemporaneous and/or compositionally similar magmatic intrusions present in the Armorican belt, and (ii) lamprophyre-type intrusive sheet swarms widespread elsewhere in the European Variscides, and more especially in the Bohemian massif.

472

473

**5.1. Emplacement of intrusive sheets in a transpressional setting**
The type of strain regime is known to be one of the governing parameters that control the trajectory and final geometry of sheet intrusions in crustal levels recording shortening (Lister and Kerr, 1991; Watanabe et al., 2002; Sibson, 2003). The transpressional setting displayed by the WCAD during the Variscan orogeny has been briefly evoked above (Section 1). It can be better defined in terms of strain regime when comparing the resulting structural patterns in both the Brest syncline and the adjoining LMD to the north. The two domains show highly contrasting tectonic styles. In the LMD, the main ductile event, i.e. the regional foliation (correlated with S₁ in the WCAD) which formed prior to exhumation path and doming processes, is characterized by planar/linear and linear (S/L- and L-types) fabrics (Authemayou et al., 2021). The corresponding NE-SW stretching lineation is associated to both wrench dextral shearing and top-to-the-NE tangential (thrust) shearing. This strain pattern has been approximated to an oblique/to/lateral flow produced by the interplay of strike-perpendicular shortening and orogen-parallel stretching. By contrast, the deformed Paleozoic series in the WCAD consistently show planar ductile fabrics (S-type, S₁/S₂ cleavages), devoid of any linear fabrics (Section 4.1), which are typically indicative, at first approximation (without any accurate strain markers), of a plane strain regime under a strike-normal shortening.

The juxtaposition of these two structurally distinct domains suggests a strain partitioning process, further associated to a major decoupling mechanism, possibly *via* the ESZ (Fig. 13a).

The strain model applied to the WCAD has major implications on the conditions of magma injection in the WAS complex. It implies that the opening of the orogen-parallel (N50-100°) magma-filled fracture network (ca. 50 %) did not occur in response to any (D₁) regional stretching. This assertion agrees with the strain/intrusion temporal pattern applied above which shows that >70% of intrusive sheets are post-D₁ magmatic features. In these conditions, two alternative driving mechanisms might be responsible for the orientation and opening of the orogen-parallel magma-filled fracture network in the WAS complex. Firstly, it has been
demonstrated above that: (i) 40% of WAS intrusions are oriented parallel to the regional S₁ cleavage and (ii) they were emplaced as either syn-S₁ (Le Roz shear zone) or post-S₁ structures. It has also been shown that 13 shallowly-dipping intrusions follow the trace of D₃ shear zones. It thus clearly appears that host-rocks tectonic anisotropies played a crucial role on the propagation of magma, whatever their (pre- or syn) origins with regards to the successive magmatic pulses. Interplays of shearing and magma injection are conceivable because of the lubricant behavior usually attributed to shear planes (Hollister and Crawford, 1986). Though no data yet exist about magma properties and ambient tectonic stresses, it can be reasonably hypothesized that magma overpressure acted as a potential driving force during magma injection, provided that: (i) it overcame the applied tectonic stresses, (ii) it was equal to or higher than the vertical stress (Goult y, 2005), and (iii) it was sufficient to counter the normal stress on planar discontinuities (Delaney et al., 1986; Meriaux and Lister, 2002). This hypothesis agrees with similar studies showing that magma pressure can promote upward migration of magma through upper crustal terrains subject to shortening (Tibaldi, 2005).

Two alternative driving mechanisms of magma injection are envisaged when considering the earliest magma-filled fracture network coeval to the S₁ event, i.e. those present in the South Chateau and Le Roz sites. At the South Chateau site, the N120-140°E-oriented intrusions, at high angle to the regional sheet swarm, likely originated as vertical tension-gashes, parallel to the shortening direction (Fig. 13a). These are the only WAS magmatic structures compatible with the regional shortening pattern and thus consistent with theoretical models (Anderson, 1951; Glazner, 1991; Hamilton, 1995; Watanabe et al., 1999). By contrast, the orogen-parallel strike of syn-S₁ intrusions in the Le Roz feeder/shear zone (Fig. 6) does not fit with theoretical predictions. From regional considerations, structural inheritance is regarded as a second possible driving mechanism for magma emplacement in one specific part of the WAS complex. On the sketch structural maps in Figs. 13a, b, the similar orientations and (dextral) kinematics
of the ESZ, Le Roz and MASZ structures suggest that they belong a common regional fault pattern. Therefore, the inherited (Early Paleozoic) origin of the ESZ (Section 1) should also apply to the Le Roz and MASZ structures. These colinear structures likely nucleated along a N70°E-oriented discontinuity and later acted as magmatic feeder zones for lamprophyric and granitic syntectonic magmas, respectively, while recording high strain dextral shearing (Fig. 13). The fact that high-strain shearing spatially coincided with the Le Roz magmatic feeder zone is thought not to be fortuitous and might be assigned to focusing of strain in a crustal portion thermally weakened by magma. Such strain/magmatism feedback interactions are routinely described in actively deforming shear zones (Druguet et al., 2008; Butler and Torvela, 2018; Alsop et al., 2021).

The results above clearly show that theoretical models about syncontractional intrusive sheets (e.g., Hubbert and Willis, 1957; Sibson, 2003; Tibaldi, 2005; Gonzales et al., 2009) only apply to a moderate amount of WAS magma-filled fractures, i.e. those dipping shallowly or oriented parallel to the shortening direction. These are the only intrusions thought to have been directly controlled by the ambient tectonic stresses. Conversely, two alternative driving forces, i.e. host-rock tectonic anisotropies ($S_1$ and $D_3$) and structural inheritance, likely operated during the emplacement of the prominent orogen-parallel WAS intrusion population. In such contexts, the mean orientation of most intrusive sheets does not pertain to a direct relation to the regional stress/strain field and thus cannot be used as a proxy for time-averaged stresses (Gudmundsson, 1995; Sato et al., 2013).

## 5.2. The WAS complex in the Armorican Variscan context

The Variscan orogeny in the Armorican massif was accompanied by a voluminous and relatively homogeneous magmatic cortege, dominated by massive granitoid intrusions. Most of them are parts of a synorogenic intrusive system formed in the time-range 320-305 Ma
(Bernard-Griffith et al., 1985; Tartèse et al., 2011; Le Gall et al., 2014) along a composite crustal-scale dextral shear zone pattern comprising of the NASZ (Saint-Renan, Huelgoat, Commana/Plouaret granites) and the SASZ (Pontivy, Lizio, Questembert granites; Fig. 13b). Younger synkinematic granitic plutons were emplaced in three various structural contexts (Fig. 13b). At ca. 300 Ma, the Quintin granite intruded along the central part of the NASZ (Peucat et al., 1984), whereas the Aber-Ildut/Brignogan granite association (Marcoux et al., 2009) was emplaced during sinistral shearing along the N70°E-oriented Porspoder-Guisseny lineament in the LMD (Chauris, 1994). The kinematic change from dextral to sinistral shearing resulted from the clockwise rotation (from NW-SE to N-S) of the applied stress field (Rolet et al., 1986). A third group of granites (Carnac, Guérande, Quiberon, Sarzeau, in Fig. 13b) were emplaced during extensional tectonics in the time-span 310-300 Ma (Le Hebel et al., 2002; Turillot et al., 2011) and thus marked fundamental geodynamical changes related to the gravitational collapse of the chain in the South Armorican domain (Gapais et al. 1993; Cagnard et al., 2004).

From the brief review above, it thus appears that the only common structural feature displayed by the WAS complex and others Armorican Variscan magmatic intrusions is the inherited origin and collinear map-trace of its inferred feeder zone (Le Roz) with those of the Huelgoat granite (MASZ; Fig. 13). These deeply-rooted discontinuities acted, possibly diachronously, as pathways for lamprophyric and granitic magmas, in response to abrupt lateral heterogeneities at depth in the mantle during the Variscan events.

But, more generally, the WAS complex radically differs from the synorogenic Armorican magmatic systems with its specific lamprophyre-type nature. Another lamprophyric network occurs in the NW part of the LMD (Fig. 13b), but it is composed of leucominettes (Caroff et al., 2015) derived from very different compositionally products and further emplaced in a quite distinct structural context. Indeed, the leucominette swarm, genetically related to the ca. 300 Ma-old Aber-Ildut granite (Caroff et al., 2015), only comprises a few dykes, approximated to
N-S tension-gashes formed parallel to the regional shortening axis (Rolet et al., 1986) during a late orogenic sinistral shearing event.

With regards to the Armorican Variscan framework, the WAS complex appears as the only lamprophyre intrusive sheet network intruded during the main shortening phase. Therefore, the nature of the causal mechanisms of such a magmatic and structural specificity should be questioned. Elements of answer are supplied from its restricted spatial distribution in the Brest syncline, with a preferred focus in the youngest (Middle/Upper Devonian) series coring its axial part (Figs. 2a, 13a). This is exemplified along the northern edge of the syncline where the Brioverian schist series occurring immediately north of the Elorn bounding shear zone, beyond the limits of the initial Paleozoic basin, are devoid of any intrusive sheets despite the proximity of the Caro potential feeder zone. The influence of a lithological/rheological contrast on magma distribution is excluded because: (i) Brioverian schists are a relatively weak material that should have rather helped magma injection, and (ii) these series are necessarily intruded at depth underneath the Paleozoic succession filling the Brest syncline. One possible mechanism, illustrated on the lithospheric-scale sketch in Fig. 15, is based on the following evidence. Since highly-displaced allochthonous units are not present in the moderately-strained Brest syncline, the map distribution of WAS intrusions roughly coincides with the main Paleozoic sedimentary depocenter. Because of the lack of any Variscan crustal-scale thrust structures, the initial (basinal) lithosphere profile must have been preserved during the orogenic events. Consequently, the necking zone of the stretched lithosphere, inferred to still underlie the Paleozoic depocenter, may have later favored ponding of melting sources directly beneath the Brest syncline. This assumption agrees with the geobarometric results of Caroff et al. (2021) which indicate an anomalously shallow paleodepth of 80 ± 5 km for the melting sources of the kersantite magmas at the base of the lithosphere.
599

**5.3. The WAS complex in the magmatic context of the European Variscides**

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Lamprophyre-type magmatism is common throughout the European Variscides, for instance in (Fig. 1a): (i) the South Bohemian massif (Zeitlhofer et al., 2014), (ii) the Black Forest (Hegner et al., 1998), (iii) the Western Swiss Alps (Monjoie et al., 2007; Filippi et al., 2019), and (iv) the Central Iberian Zone (Scarrow et al., 2011). Most of these complexes display a wide spectrum of magmatic affinities, typically indicative of both orogenic and anorogenic systems. They are generally related to wrench tectonics in either transpressional or transtensional settings (for review see Scarrow et al., 2011). The great variety of lamprophyric magmas injected at various stages of the Variscan orogeny is particularly well expressed in the SE Bohemian massif (Fig. 1b). Because most of the corresponding published works are mainly devoted to strictly magmatic (geochemical and radiometric) issues (for review see Krmicek et al., 2020), direct comparisons with the structural results obtained here about the WAS complex are made difficult. Nevertheless, correlations can be attempted between successive generations of lamprophyre intrusions reported in both the NW Armorican and SE Bohemian segments, though significant differences exist in terms of temporal patterns and structural significance in the orogen history. In the Moldanubian zone (SE Bohemian massif, Figs. 1a, b), a hiatus of ca. 10 Ma. separates the emplacement of an early intrusive sheet network showing orogenic magmatic affinities (Neubauer et al., 2003; Kubinova et al., 2017), possibly comparable to the WAS complex, and younger lamprophyre intrusions typically displaying anorogenic affinities (Fig. 1c; Neubauer et al., 2003). By contrast, a continuum seems to exist in the NW Armorican lamprophyre system between the injection of the WAS complex during the main shortening event and the ca. 300 Ma-old leucominette dyke swarm in the LMD which still displays orogenic magmatic affinities (Caroff et al., 2015), in agreement with its emplacement in a sinistral transpressional setting at a later orogenic stage.
At a larger scale, the genetic links evoked above between the Variscan lamprophyre-type magmatism and strike-slip tectonics are also mentioned in the compiled work of Scarrow et al. (2011) about older orogens, Archaean (Canada) or Caledonian in age (E Ireland, NW England), as well as about younger Cenozoic orogens (NW China, Tibet, Balkans). These links have been variously interpreted and, for example, in the model proposed by Vaughan and Scarrow (2003) K-rich mantle-derived (low-degree) melts tend to promote focusing of strain along steep discontinuities prone to be subsequently reactivated as wrench shear zones, depending on the stress field. At a first order, this model should apply to the WAS complex, given the dextral shear zone nature of the Le Roz feeder zone, but with caution because of the role assigned to structural inheritance in our model.

633

## 6. Conclusions

- Our structural study is devoted to a lamprophyric sheet swarm cutting through low-grade and moderately-deformed Paleozoic metasedimentary host-rocks occurring in a Variscan transpressional setting at the western termination of the Central Armorican Domain, western France.

- The so-called West Armorican Swarm (WAS) is composed of nearly 100 individual intrusive sheets (mostly kersantite), identified on a reference coastal cross-section, and analyzed in terms of geometry, propagation and timing of emplacement with respect to the strained host-rocks.

- Strain/intrusion mutual relationships indicate that magma injection took place during the regional shortening, *via* multiple pulses which alternated with three strain events ($D_1$, $D_2$, $D_3$).
646 - Most of individual sheets are: (i) <5 m in width, (ii) moderately- to steeply-inclined (90% >30°), and (iii) preferentially oriented N70-90°E, parallel to the trace of the regional S₁ cleavage, i.e. nearly normal to the direction of shortening.
647
648

649 - These structural features do not match with theoretical models. They are satisfactorily integrated in a tectonic model which favours the influence of both host-rock planar anisotropies (S₁ cleavage and D₃ shear structures) and structural inheritance on the propagation of magma in the WAS complex. The role of the ambient stress field is only hypothesized about a few WAS intrusions, parallel to the shortening direction.
650
651
652
653

654 - The WAS intrusions generally escaped deformation. Most of strained (boudinaged) 655 intrusive sheets occur in an inferred magmatic feeder zone (Le Roz), further reactivated as a 656 dextral shear zone.
657

657 - At the scale of the European Variscan belt, syncontractional lamprophyre complexes 658 comparable to the WAS system are only reported in the SE Bohemian massif.
659

659 - Considering the WAS complex as a reference example for lamprophyre magmas 660 emplaced in a shortened mid/upper crust requires the following conditions: (i) a subduction 661 zone and associated mantle-derived melts, (ii) a previous long-lived and subsiding sedimentary 662 basin (with its thick pile of sediments) underlain by a lithospheric necking zone, still preserved 663 during the orogenic stage, and which will later promote ponding of the melting material. The 664 modest dimensions of the initial basin, mirrored at depth by those of the necking zone, accounts 665 for the spatially-restricted distribution of the syncollisional magmatic products. Lastly, (iii) a 666 steep network of crustal-scale discontinuities, subsequently recording dextral shearing in the 667 present case (Variscan transpressional setting), act as propagation pathways (feeder zones) for 668 the deep magma up to shallower crustal levels.
669
670 Declaration of Competing Interest

671 No competing interests.

672

673 Data availability

674 Data will be made available on request.

675

676 Acknowledgements

677 Funding of our work was mostly supplied by the Parc Naturel Regional Armorique in the
678 framework of the (UNESCO) « Armorique Geopark” project conducted by Jérémie
679 Bourdoulous and Noémie Courant who are warmly thanked. Our colleagues Jean-Alix Barrat,
680 Christine Authemayou, Gilles Chazot and Laurent Geoffroy (Geo-Ocean, UMR 6538), as well
681 as Mohamed Ahmed Daoud (CERD Djibouti), are acknowledged for fruitful discussions in the
682 field. Dr. José Ramon Catalàn (Salamanca univ.) provided useful comments that have been
683 appreciated. Kelly Cayocca (Troglogites Monts d’Arrée) is thanked for “polishing” the English
684 form of parts of the text. We gratefully acknowledge Dr. S. Angiboust (Editor) and the
685 reviewers M. Faure and one anonymous, with special thanks to D. Zanoni for his thorough
686 examination of the paper and his constructive comments.

687

688 Supplementary data

689 Supplementary data to this article can be found online...

690

691 CRediT authorship contribution statement

692 Bernard Le Gall: Conceptualization, Methodology, Investigation, Writing, Visualization.

693 Martial Caroff: Investigation, Formal analysis, Validation, Funding acquisition.

694

695

696 References
697 Alsop, G., Strachan, R., Holdsworth, R., Burns, I., 2021. Geometry of folded and boudinaged pegmatite veins emplaced within a strike-slip shear zone: a case study from the Caledonian orogen, northern Scotland. J. Struct. Geol. 142, 104233.

700 Anderson, E., 1951. The dynamics of faulting and dyke formation with applications to Britain. Oliver & Boyd, Edinburgh.

702 Authemayou, C., Le Gall, B., Caroff, M., Bussien-Grosjean, D., 2019. Wrench-related dome formation and subsequent orogenic syntax bending in a hot orogen (Variscan Ibero-Armican Arc, the Ouessant Island, France). Tectonics 38, https://dx.doi.org/10.1029/2018TC005189.

706 Babin, C., Darboux, J.R., Duée, G., Gravelle, M., Morzadec, P., Plusquellec, Y., Thonon, P., 1975. Tectoniques tangentielles et tectoniques superposées dans le Dévonien de la Rade de Brest. C. R. Acad., Sc., Paris 280, 259-262.

709 Babin, C., Darboux, J.R. (coordinators), 1982. Geological map of France, sheet Le Faou (275). BRGM, Orléans, scale 1:50,000.

711 Balé, P., Brun, J.-P., 1986. Les complexes métamorphiques du Léon (NW Brittany): un segment du domaine éohercynien sud-armoricain translaté au Dévonien. Bull. Soc. Geol. Fr. 8(2), 471-477.

714 Ballèvre, M., Bosse, V., Ducassou, C., Pitra, P., 2009. Paleozoic history of the Armorican Massif. Models for the tectonic evolution of the suture zones. C. R. Geosci. 341, 174-201.

716 Ballouard, C., 2016. Origine, évolution et exhumation des leucogranites peralumineux de la chaîne hercynienne armoricaine : implication sur la métallogénie de l'uranium. Ph.D. thesis. Rennes1 univ. France.

719 Bernard-Griffiths, J., Peucat, J., Sheppard, S., Vidal, P., 1985. Petrogenesis of Hercynian leucogranites from South Armorican massif. Contributions of REE and isotopic (Sr, Nd,
721 Pb, O) geochemical data to the study of source rock characteristics and ages. Earth Planet.

722 Sci. Lett. 74, 235-250.

723 Bertelsen, H., Rogers, B., Galland, O., Dumazer, G., Bennani, A., 2018. Laboratory modelling
724 of coeval brittle and ductile deformation during magma emplacement into viscoelastic
725 rocks. Front. Earth Sci. 6. https://doi.org/10.3389/feart.2018.00199.

726 Butler, R., Torvela, T., 2018. The competition between rates of deformation and solidification
727 in synkinematic granitic intrusions: resolving the pegmatite paradox. J. Struct. Geol. 117,
728 1-13.

729 Cagnard, F., Gapais, D., Brun, J.-P., Gumiaux, C., van den Driessche, J., 2004. Late pervasive
730 crustal-scale extension in the south Armorican Hercynian belt (Vendée, France). J. Struct.
731 Geol. 26, 435-449.

732 Caroff, M., Barrat, J.-A., Le Gall, B., 2021. Kersantites and associated intrusives from the type
733 locality (Kersanton), Variscan belt of Western Armorica (France). Gondwana Res. 98, 46-
734 62.

735 Caroff, M., Labry, C., Le Gall, B., Authemayou, C., Bussien Grosjean, D., Guillong, M., 2015.
736 Petrogenesis of late-Variscan high-K alkali-calcic granitoids and calc-alkalic
737 lamprophyres: The Aber-Ildut/North-Ouessant complex, Armorican Massif, France. Lithos
738 238, 140-155. http://dx.doi.org/10.1016/j.lithos.2015.09.025.

739 Caroff, M., Le Gall, B., Authemayou, C., 2020. How does a monzogranite turn into a
740 trachydacitic extrusion mantled by basinal volcaniclastics and peperites? The case of South
741 Ouessant, Armorican Variscides (France). J. Geol. Soc., London doi.org/10.1144/jgs2020-
742 060.

743 Caroff, M., Le Gall, B., Authemayou, C., Bussien Grosjean, D., Labry, C., Guillong, M. 2016.
744 Relations between basalts and adakite-felsic intrusive bodies in a soft-substrate
745 environment: the South Ouessant Visean basin in the Variscan belt, Armorican Massif,
746 France. Can. J. Earth Sci., 53, 441-456, https://doi.org/10.1139/cjes-2015-0230.

747 Castaing, C. et al., 1987. Geological map of France, sheet Huelgoat (276). BRGM, Orléans,
748 scale 1: 50,000.

749 Chauris, L., 1994. Geological map of France, sheet Plouarzel/île d'Ouessant (295), BRGM,
750 Orléans, scale 1: 50,000.

751 Chauris, L., Plusquellec, Y. et al., 1980. Geological map of France, sheet Brest (274). BRGM,
752 Orléans, scale 1: 50,000.

753 Darboux, J.-R., Le Gall, B. 1988. Les Montagnes Noires : cisaillement bordier méridional du
754 bassin carbonifère de Châteaulin (Massif Armoricain, France). Caractéristiques
755 structurales et métamorphiques. Geodinamica Acta 2(3), 121-133.

756 Darboux, J.-R., Marcoux, E., Hallégouët, B., Lebret, P., Sorel, P., 2010. Geological map of
757 France, sheet Landerneau (239). BRGM, Orléans, scale 1: 50,000.

758 Delaney, P., Pollard, D., Ziony, J., McKee, E., 1986. Field relations between dikes and joints:
759 emplacement processes and paleostress analysis. J. Geophys. Res. Solid Earth 91 (B5),
760 4920-4938.

761 Doubinger, J., Pelhate, A., 1976. Nouvelles observations sur l'âge des schistes de Châteaulin
762 (Massif armoricain). C. R. Acad. Sci., Paris 283, 467.

763 Druguet, E., Czeck, D., Carreras, J., Castano, L., 2008. Emplacement and deformation features
764 of syntectonic leucocratic veins from Rainy Lake zone (Western Superior Province,
765 Canada). Precambrian Res. 163 (3-4), 384-400.

766 Faure, M., Bé Mézème, E., Cocherie, A., Rossi, P., Chemenda, A., Boutelier, D., 2008.
767 Devonian geodynamic evolution of the Variscan Belt, insights from the French Massif
768 Central and Massif Armoricain. Tectonics 27, doi.org/10.1029/2007 TC002115.
769 Faure, M., Sommers, C., Melleton, J., Cocherie, A., Lautout, O., 2010. The Léon domain (French Massif armoricain): a westward extension of the Mid-German Crystalline Rise?
770
771 Structural and geochronological insights. Int. J. Earth Sci. 99, 65-81.

772 Ferré, E., Galland, O., Montanari, D., Kalakay, T., 2012. Granite magma migration and emplacement along thrusts. Int. J. Earth Sci. 1-16. doi: 10.1007/s00531-012-0747-6.

774 Filippi, M., Zanoni, D., Gosso, G., Lardeaux, J.-M., Verati, C., Spalla, M.I., 2019. Structure of lamprophyres: a discriminant marker for Variscan and Alpine tectonics in the Argentera-Mercantour Massif, Maritime Alps. Bull. Soc. Géol. France 190, 1-16.
776
777 https://doi.org/10.1051/bsgf/2019014

778 Foster, D., Schafer, C., Fanning, M., Hyndman, D., 2001. Relationships between crustal partial melting, plutonism, orogeny, and exhumation: Idaho-Bitterroot batholith, Tectonophysics 342, 313-350.

781 Galland, O., de Bremond d'Ars, J., Cobbold, P., Hallot, E., 2003. Physical models of magmatic intrusion during thrusting. Terra Nova 15, 405-409.

783 Galland, O., Cobbold, P., de Bremond d'Ars, J., Hallot, E., 2007a. Rise and emplacement of magma during horizontal shortening of the brittle crust: Insights from experimental modeling. J. Geophys. Res. 112, B06402, doi:10.1029/2006JB004604.

786 Galland, O., Hallot, E., Cobbold, P., Ruffet, G., de Bremond d'Ars, J., 2007b. Volcanism in a compressional Andean setting: A structural and geochronological study of Tromen volcano (Neuquén province, Argentina). Tectonics 26, TC4010, doi:10.1029/2006TC002011.

789 Gapais, D., Le Corre, C., 1980. Is the Hercynian belt of Brittany a major shear zone? Nature 288, 574-576.

791 Gapais, D., Lagarde, J., Le Corre, C., Audren, C., Jegouzo, P., Casas Sainz, A., van den Driessche, J., 1993. La zone de cisaillement de Quiberon : témoin d'extension de la chaîne varisque en Bretagne méridionale au Carbonifère. C. R. Acad. Sci., Paris 316, 1123-1129.
794 Glazner, A., 1991. Plutonism, oblique subduction, and continental growth: An example from
795 the Mesozoic of California, Geology 19, 784-786.

796 Gonzalez, G., Cembrano, J., Aron, F., Veloso, E., Shyu, J., 2009. Coeval compressional
797 deformation and volcanism in the central Andes, case studies from northern Chile (23°S-
798 24°S). Tectonics 28, 6, TC6003

799 Goult y, N., 2005. Emplacement mechanism of the Great Whin and Midland Valley dolerite
800 sills. J. Geol., Soc., London 162, 1047-1056.

801 Gudmundsson, A., 1995. Infrastructure and mechanics of volcanic systems in Iceland. J. Volc.
802 Geother. Res. 64/1-2, 1-22.

803 Gumiaux, C., Judenherc, S., Brun, J.-P., Gapais, D., Granet, M., Poupinet, G., 2004. Restoration
804 of lithosphere-scale wrenching from integrated structural and tomographic data (Hercynian
805 belt of western France). Geology 32, 333-336.

806 Hamilton, W., 1995. Subduction systems and magmatism. In: J., Smellie (ed.), Volcanism
807 associated with extension at consuming plate margins. Geol. Soc. Lond., Spec. Pub. 3-28.

808 Hegner, E., Kolbl-Ebert, M., Loeschke, J., 1998. Post-collisional Variscan lamprophyres (Black
809 Forest, Germany): $ ^{40}Ar / ^{39}Ar $ phlogopite dating, Nd, Pb, Sr isotope, and trace element
810 characteristics. Lithos 45, 395-411.

811 Hollister, L., Crawford, M., 1986. Melt-enhanced deformation: a major tectonic process.
812 Geology 14, 558-561.

813 Hubbert, M., Willis, D., 1957. Mechanics of hydraulic fracturing. In: Hubbert, M., (ed),
814 Structural Geology 175-190, Macmillan, New York.

815 Jones, K., 1994. Progressive metamorphism in a crustal-scale shear zone: an example from the
816 Leon region, NW Brittany, France. J. Metam. Geol. 12, 69-88.
817 Kjøll, H., Andersen, T., Corfu, F., Labrousse, L., Tegner, C., Abdelmalak, M., Planke, S., 2019.

818 Timing of break-up and thermal evolution of a pre-Caledonian Neoproterozoic exhumed

819 magma-rich rifted margin. Tectonics doi.org/10.1029/2018TC005375.

820 Krmicek, L., Romer, R., Timmerman, M., Ulrych, J., Glodny, J., Prichystal, A., Sudo, M., 2020.

821 Long-lasting (65 Ma) regionally contrasting late- to post-orogenic Variscan mantle-derived

822 potassic magmatism in the Bohemian massif. J. Pet. Geol. 61, 7, doi:

823 10.1093/petrology/egaa072.

824 Kroner, U., Romer, R., 2013. Two plates - Many subduction zones: The Variscan orogeny

825 reconsidered. Gondwana Res. 24, 298-329.

826 Kubinova, S., Faryad, S., Verner, K., Schmitz, M., Holub, F. 2017. Ultrapotassic dykes in the

827 Moldanubian Zone and their significance for understanding of the post-collisional mantle

828 dynamics during Variscan orogeny in the Bohemian Massif. Lithos 272-273, 205-221.

829 Le Gall, B., Authemayou, C., Ehrhold, A., Paquette, J.-L., Bussien, D., Chazot, G., et al., 2014.

830 LiDAR offshore structural mapping and U/Pb zircon/monazite dating of Variscan strain in

831 the Léon metamorphic domain, NW Brittany. Tectonophysics 630, 236-250.

832 https://doi.org/10.1016/j.tecto.2014.05.026.

833 Le Gall, B., Authemayou, C., Graindorge, D., Duperret, A., Kaci, T., Ehrhold, A., Schmitt, T.

834 2021. Status of basement terranes in orogens : Insights from the Cadomian Domain of the

835 North Armorican Variscides, Western France. Tectonics 40(5), doi:

836 10.1029/2020TC006578.

837 Le Gall, B., Loboziak, S., Le Herissé, A., 1992. Le flanc sud du synclinorium carbonifère

838 de Châteaulin (Massif Armoricain, France) : une bordure de bassin reactivée en contexte

839 decro-chevauchant. Bull., Soc., geol., Fr. 163, 1, 13-26.

840 Le Hebel, F., Vidal, O., Kienast, J., Gapais, D., 2002. Les porphyroïdes de Bretagne

841 méridionale : une unité de HP-BT dans la chaîne hercynienne. C. R. Géosci. 334, 205-211.
842 Lister, J., Kerr, R., 1991. Fluid-mechanical models of crack propagation and their application
to magma transport in dykes. J. Geophys. Res. 96, 10049-10077.

844 Marcoux, E., Cocherie, A., Ruffet, G., Darboux, J.-R., Guerrot, C., 2009. Géochronologie
revisitée du dôme du Léon (Massif armoricain, France). Géol. Fr. 1, 19-40.

846 Mazzarini, F., Musumeci, G., Montanari, D., Corti, G., 2010. Relations between deformation
and upper crustal magma emplacement in laboratory physical models. Tectonophysics 484,
13-146.

849 Menand, T., Daniels, K., Benghiat, P., 2010. Dyke propagation and sill formation in a
compressive tectonic environment. J. Geophys. Res. 115, doi:10.1029/2009JB006791.

851 Meriaux, C., Lister, J., 2002. Calculation of dike trajectories from volcanic centers. J. Geophys.
Res. 107, B4, doi:10.1029/2001JB000436.

853 Monjoie, P., Bussy, F., Schaltegger, U., Mulch, A., Lapierre, H., Pfeifer, H., 2007. Contrasting
magma types and timing of intrusion in the Permian layered mafic complex of Mont Collon
(western Alps, Valais, Switzerland): Evidence from U/Pb zircon and 40Ar/39Ar amphibole
dating. Swiss J. Geosci. 100, 125-135.

857 Montanari, D., Corti, G., Sani, F., Del Ventisette, C., Bonini, M., Moratti, G., 2010a.
Experimental investigation on granite emplacement during shortening. Tectonophysics
484, 1-4, 147-155.

860 Montanari, D., Corti, G., Simakin, A., 2010b. Magma chambers and localization of deformation
during thrusting. Terra Nova 22, 390-395.

862 Neubauer, F., Dallmeyer, D., Fritz, H., 2003. Chronological constraints of late-and post-
orogenic emplacement of lamprophyre dykes in the southeastern Bohemian Massif,
Austria. Swiss Bull. Mineral. Petrol. 83, 317-330.

865 Paradis, S., Velde, B., Nicot, E., 1983. Chloritoid-pyrophyllite-rectorite facies rocks from
Brittany, France. Contrib. Min. Pet. 83, 342-347.
867 Pasquare, F., Tibaldi, A., 2007. Structure of a sheet-laccolith system revealing the interplay
868 between tectonic and magma stresses at Stardalur Volcano, Iceland. J. Volc. Geotherm.
869 Res. 161, 131-150.

870 Peucat, J.-J., Auvray B., Hirbec Y., Calvez J.-Y., 1984. Granites et cisaillements hercyniens
dans le Nord du Massif Armoricain : géochronologie Rb-Sr, Bull. Soc. géol. France 7, 26,
871 6, 1365-1373.

873 Plusquellec, Y., Rolet, J., Darboux, J.-R. et al., 1999. Geological map of France, sheet
874 Châteaulin (310). BRGM, Orléans, scale 1: 50,000.

875 Pons, J., Barbey, P., Dupuis, D., Leger, J., 1995. Mechanisms of pluton emplacement and
876 structural evolution of a 2.1 Ga juvenile continental crust: the Birimian of southwestern
877 Niger. Precambrian Res. 70, 281-301.

878 Rolet, J., Thonon, P., 1978. La semelle d'un charriage hercynien majeur effondrée par un réseau
de fractures en régime coulissant dextre : sa mise en évidence grâce aux marqueurs
879 filoniens de la Rade de Brest (Massif armoricain). C. R. Acad. Sci., Paris 287, 1099-1102.
880

881 Rolet, J., Le Gall, B., Darboux, J.-R., Thonon, P., Gravelle, M., 1986. L'évolution
géodynamique dévono-carbonifère de l'extrémité occidentale de la chaîne hercynienne
882 d'Europe sur le transect Armorique-Cornwall. Bull. Soc. Geol. Fr. 8 (2), 43-54.
883

884 Scarrow, J., Molina, J., Bea, F., Montero, P., Vaughan, A., 2011. Lamprophyre dikes as markers
of late orogenic transtension timing and kinematics: a case study from the Central Iberian
885 Zone. Tectonics 30, doi:10.1029/2010TC002755.

887 Sato K, Yamaji, A., Tonai, S., 2013. Parametric and non-parametric statistical approaches to
the determination of paleostress from dilatant fractures: application to an Early Miocene
888 dike swarm in central Japan. Tectonophysics 588, 69-81.

890 Schofield, N., Brown, D., Magee, C., Stevenson, C. 2012. Sill morphology and comparison of
brittle and non-brittle emplacement mechanisms. J. Geol. Soc., London 169, 127-141.
892 Schulz, B., 2013. Monazite EMP-Th-U/Pb age pattern in Variscan metamorphic units in the
893 Armorican Massif (Brittany, France). German J. Geosci. 164, 313-335.

894 Sibson, R., 2003. Brittle-failure controls on maximum sustainable overpressure in different
tectonic regimes. Am. Assoc. Petrol. Geol. Bull. 87(6), 901-908.

896 Takada, A., 1988. Subvolcanic structure of the central dike swarm associated with the ring
complexes in the Shitara district, central Japon. Bull. Volc. 50, 106-118.

898 Tartèse, R., Boulvais, P., Poujol, M., Chevalier, T., Paquette, J.-L., Ireland, T., Deloule, E.,
899 2011. Mylonites of the South Armorican Shear Zone: Insights for crustal-scale flow and
water-rock interaction processes. J. Geodyn. doi:10.1016/j.jog.2011.05.003.
900

901 Tibaldi, A., 2005. Volcanism in compressional tectonic settings: Is it possible? Geophys. Res.
902 Lett. 32, doi:10.1029/2004GL021798.

903 Tibaldi, A., Bonali, F., Corazzato, C., 2017. Structural control on volcanoes and magma paths
904 from local- to orogen-scale: The central Andes case. Tectonophysics 699, 16-41.

905 Thonon, P., Rolet, J., 1982. Magmatisme et tectonique en Domaine Centre Armoricain
occidental (Finistère). 107 Cong. Nat. Soc. Sav., Brest 355-366.

907 Turrillot, P., Augier, R., Monié, P., Faure, M., 2011. Late orogenic exhumation of the Variscan
908 high-grade units (South Armorican Domain, western France), combined structural and
40Ar/39Ar constraints. Tectonics 30, http://dx.doi.org/10.1029/2010TC002788.
909

910 van Noorden, M., Sintubin, M., Darboux, J.-R., 2007. Strain partitioning in a slate belt:
911 Evidence from the early Variscan Monts d'Arrée slate belt (Brittany, France). J. Struct.
912 Geol 29, 837-849.

913 Vaughan, A., Scarrow, J., 2003. K-rich mantle metasomatism control of localization and
initiation of lithospheric strike-slip faulting. Terra Nova 15, 163-169.

915 von Raumer, J., Finger, F., Veselá, P., Stampfli, G., 2014. Durbachites-Vaugnerites - a
geodynamic marker in the central European Variscan orogen. Terra Nova 26, 85-95.
917 Walker, G., 1999. Volcanic rift zones and their intrusion swarms. J. Volc. Geotherm. Res. 94,
918 21-34.

919 Watanabe, T., Koyaguchi, T., Seno, T., 1999. Tectonic stress controls on ascent and
920 emplacement of magmas. J. Volc. Geotherm. Res. 91, 65-78.

921 Watanabe, T., Masuyama, T., Nagaoka, K., Tahara, T., 2002. Analog experiments on magma-
922 filled cracks: competition between external stresses and internal pressure, Earth Planets
923 Space 54, 1247-1261.

924 Zeitlhofer, H., Schneider, D., Grasemann, B., Konstantin, P., Thöni, M., 2014. Polyphase
tectonics and late Variscan extension in Austria (Moldanubian Zone, Strudengau area). Int.
925 J. Earth Sc. 103, 83-102.

927 Zeitlhofer, H., Grasemann, B., Petrakakis, K., 2016. Variscan potassic magmatism of
928 durbachitic affinity at the southern end of the Bohemian Massif (Lower Austria). Int. J.
929 Earth Sc. 105, 1175-1197.

930

931 **Captions**

932 Fig. 1. (2 columns)

933 Geological setting. (a). The Armorican massif in a sketch map of the European Variscides,
934 location of Figs. 1b and 1d. (b) Simplified magmatic map of the Bohemian massif showing the
935 widespread distribution of lamprophyric complexes, modified from Krmicek et al. (2020). (c).
936 Initial εNd/age diagram showing the orogenic vs anorogenic origin of the lamprophyric
937 magmatism in the Bohemian and Armorican Variscan massifs, modified from Krmicek et al.
938 (2020). (d) The main structural domains in the Armorican Variscides, Western France. CAD.,
939 Centre Armorican domain; CS., Châteaulin syncline; ESZ., MASZ., MNSZ., NASZ., SASZ.,
940 Elorn, Monts d'Arrée, Montagnes Noires, North Armorican, South Armorican shear zone,
941 respectively; LMD., Leon metamorphic domain, SAD., South Armorican domain. The N70°-
942 oriented LMD dome structure is not drawn. (e). Schematic structural cross-section at the
943 transition from: (i) medium/high-grade metamorphic rocks in the southern flank of the LMD
944 dome and (ii) low-grade metamorphic rocks in the WCAD. The continuity of the tectonic and
945 metamorphic patterns is only disrupted by a network of discrete dextral shear zones
946 (explanations and corresponding references in the text). ESZ, NASZ and PSZ., Elorn, North
947 Armorican and Pierres Noires shear zones, respectively, location in Fig. 1d.

948

Fig. 2. (2 columns)

950 Structural sketch map of the Brest syncline. Intrusions are shown as point-like features, from
951 1: 50 000 geological maps (references in the text). C., Château headland; Ca., Caro; ESZ., Elorn
952 shear zone; FT., transverse fault; G. Gluziau; I., Illien; IL., Ile Longue; K., Kerascoet; L.,
953 Larmor; LRSZ., Le Roz shear zone; NASZ., North Armorican shear zone; P., Porsguen. RT.,
954 Rozegat thrust, location in Fig. 1d. (b). Lithostratigraphy of the metasedimentary formations in
955 the Brest syncline. (c). Structural cross-section in the Brest syncline. Specific stratigraphic
956 formations are drawn as structural markers. Their stratigraphic position is shown in Fig. 2b.
957 The WAS intrusive sheets, as well as the structures at depth, are not drawn.

958

Fig. 3. (2 columns)

960 Variscan ductile strain recorded by Paleozoic metasedimentary rocks in the Brest syncline. (a).
961 An overturned syncleavage fold (D₁) in calcareous alternations (Le Faou Fm., stratigraphic
962 position in Fig. 2b) exposed along the Larmor section (S₀, bedding; S₁, cleavage; location in
963 Figs. 2a, c). (b). Stereographic diagram of the strike of the S₁ cleavage in the Northern unit (in
964 %). (c). An overturned syncleavage fold (D₁) in the Southern unit (Traonliors Fm., stratigraphic
965 position in Fig. 2b). Porsguen area; location in Figs. 2a, c. (d). Stereographic diagrams of the
966 S₁ cleavage orientation in the Southern unit (equal area, lower hemisphere, Wulff projection).
967 (d1). Poles to S₁ cleavage. (d2). Density contours of poles to S₁ cleavage. The best-fit great
968 circle (dashed line) drawn by the contoured areas demonstrates folding of the S₁ cleavage by a
969 second generation of fold (D₂) displaying a mean B₂ axis plunging 30° to N240°E. (e).
970 Stereographic diagram of the strike of the S₁ cleavage planes in the Southern unit (in %). (f). A
971 shallowly-dipping shear (thrust) structure cutting through folded schist series in the Southern
972 unit. Kerlorziou section; location in Fig. 2a, Supplementary Material Table S1).

973

Fig. 4. (2 columns)

975 Spatial distribution, orientation and width of the WAS intrusion network. (a). Histogram
976 showing the amount of intrusions *per* km on the 20 km-long reference cross-section. (b, c, d).
977 Stereographic diagrams of the intrusion strike for the total population (b), in the Northern unit
978 (c) and in the Southern unit (d). The results are shown as %. (e). Histogram showing frequency-
979 distributions of dip angles for all the intrusive sheets measured in the WAS. (f). Histogram
showing frequency-distributions of widths for all the intrusive sheets measured in the WAS (see Supplementary Material Table S1).

Fig. 5. (2 columns)

Remarkable intrusive patterns in the WAS magmatic complex. (a). Kersantite ring-sheet structures observed in the South Château site. Note that the trace of the S₁ cleavage follows the margins of (i) the elliptical ring-sheets and (ii) the submeridian intrusions further east. (b, c). Transition patterns between steeply- and shallowly-inclined intrusion sheets in the South Château (b) and Caro (c) sites, location in Fig. 2a.

Fig. 6. (2 columns)

Main tectonic and magmatic structures in the Le Roz dextral shear zone. (a). Structural map showing: (i) two contrasted packages juxtaposed along an (inferred) thrust structure, and (ii) the intricate and lense-shaped geometry of most kersantite sheets in the lower (footwall) package. (b). Along-strike section showing a shallow microgranodiorite SIS thrusted over Upper Devonian schist series enclosing remnants of a kersantite sheet. (c). Structural cross-section through the highly-strained schist series (lower package) and its closely-spaced kersantite sheet network. (d). A moderately-inclined kersantite sheet (K₆₃) totally embedded in highly-sheared schists, but still displaying a well-preserved cooling joint pattern. Hammer (encircled) for scale. (e). Blocks of kersantite in a tectonic schist breccia. Hammer (encircled) for scale, location in Fig. 6c. See Supplementary Material Table S1.

Fig. 7. (2 columns)

Main tectonic and magmatic structures in the Ile Longue shear zone. (a). Sketch structural map showing two tectonic packages juxtaposed along a top-to-the-SW shear thrust contact, modified from Chauris et al. (1980). (b). Detailed structural map of the lower package showing the dismembered geometry of microgranodiorite sheets in a dextral ductile shear zone. (c). Cross-section showing the structural contrast between the two superposed packages. Trace of the section in Fig. 7a, completed and modified from Thonon and Rolet (1982).

Fig. 8. (1.5 column)

Relative chronology of strain and magmatic events in the Brest syncline. (a) Conceptual model illustrating the various ways host-rocks anisotropies have controlled the trajectory of magma in the WAS (Variscan) complex. No scale. See text for explanations. (b) Various intrusion
networks alternating with three strain events (D₁, D₂, D₃). Kersantite and microgranodiorite undifferentiated). Black and white arrows correspond to fully and partly bracketed intrusive sheet populations, respectively. See Supplementary Material Table S1.

Fig. 9. (2 columns)

Tectonic and magmatic structures in the Château headland site, location in Fig. 2a. (a). Structural map showing two generations of kersantite sheets intruding Late Devonian schist series (Traonliors Fm.). (b). Structural cross-section illustrating two kersantite sheets dipping parallel with the S₁ cleavage, in a steep (HIS K₅₁) or shallow (SIS K₅₃) attitude. The SIS K₅₃ is assumed to post-date both the D₂ fold structure and the HIS K₅₁ in the lower package.

Fig. 10. (2 columns)

Various types of temporal relationships between strain and magma emplacement. (a). Detailed map of D₂ folds (schists/sandstones alternations in the Troaon Group) disrupted (post-dated) by three moderately-inclined kersantite sheets (K₈₂-₈₄). Yelen site. (b). A thick microgranodiorite sheet (µGD₃₃) post-dated the D₁ event (S₁ enclaves) and is involved in (and pre-dated by) a top-to-the-SE thrust structure (D₃). Squiffiec site. (c). A 3 m-thick kersantite sheet (K₂₃) injected along a shallowly-inclined thrust fault cutting through steeply-dipping quartzitic strata (D₁ fold) of the Landevennec Fm. (d). Zoomed view showing an apophysis of the main sheet (upper margin) locally disrupted by shear surfaces. Intrusion drawn in blue, hammer (encircled) for scale. Larmor site, location in Fig. 2a.

Fig. 11. (1 column)

Structural evidence illustrating the emplacement of WAS intrusive sheets during shortening. (a, b). A shallowly-inclined kersantite sheet (K₈₇) in the Kerascoet site, (a) involved in (and pre-dating) a top-to-the-north (D₃) shear thrust and (b) enclosing a schist enclave (S₁) and thus post-dating the main ductile strain event (D₁), location in Fig. 2a. (c) Folded magmatic layering in a thin microgranodiorite sheet (µGD₁₇) injected along a top-to-the-NW shear thrust (D₃). Caro section, location in Figs. 2a, 5c. These timing constraints are consistent with those applied to the kersantite sheet K₂₃ (Larmor section) and the microgranodiorite µGD₃₃ (Squiffiec section), location in Fig. 10.

Fig. 12. (1 column)
Structural evidence supporting the rigid/brittle mechanical behavior of the magmatic (sheets) and host-rock material during shortening. (a). A crack/fracture network formed by brittle failure of the schist host-rocks (Le Faou Fm.) ahead of a kersantite propagating tip (K24). Hammer for scale. Larmor site. (b). A kersantite HIS, associated to the ring-sheet complex in the South Château site (K56), displaying a wedge-shaped tip in Upper Devonian schist host-rocks (Porsguen Fm.). Hammer for scale. (c). Narrow shear zones (top-to-the-east and post-dating the intrusion emplacement) along the margins of a kersantite SIS (K5). Note the still preserved vertical cooling joint pattern. Hammer for scale. Ilien ar Gwen site, location in Fig. 2a, and Supplementary Material Fig. S1.

Fig. 13. (2 columns)

The WAS lamprophyric complex in the Armorican Variscan framework. (a). Sketch structural map showing: (i) the contrasted strain regimes in the LMD and WCAD (decoupling and strain partitioning mechanisms are inferred) and (ii) the main magmatic features in the WAS complex (see text for explanations). The map-trace of WAS intrusions is only indicative with the aim to illustrate their parallelism (for most of them) with the regional cleavage (S1). Spacing of the purple lines (intrusions) roughly reflects the density of the point-like features in the map in Fig. 2a. Same abbreviations as in Fig. 2a. (b) Spatial distribution of Variscan magmatic rocks in the Armorican belt. AI., Aber-Ildut; B., Brignogan; BS., Brest syncline; Ca., Carnac; CAD., Centre Armorican Domain; Co., Commana; G., Guérande; H., Huelgoat; LMD., Leon metamorphic domain; Pm., Ploumanac'h; Po., Pontivy; Q., Quintin; S., Sarzeau; SR., Saint-Renan; ESZ., MASZ., MNSZ., NASZ., PGSZ., SASZ., Elorn, Monts d'Arrée, Montagnes Noires, North Armorican, Porspoder-Guissény, South Armorican shear zones, respectively.

Fig. 14. (2 columns)

Dip intrusion distribution plotted against strikes. The two grey zones correspond to intrusive sheets that contradict theoretical models in terms of: (i) high (HIS) to moderate (MIS) dipping attitudes (dark grey) and (ii) orientation with respect to the direction of the regional shortening (Z) (light grey).

Fig. 15. (2 columns)

The WAS magmatic system in the lithospheric framework of the West Central Armorican Domain. The paleodepths of the crustal magmatic chamber and the melting sources in the necking lithosphere zone are from Caroff et al. (2021). The dashed line has no geological
1081 meaning (base of the section). CF., Caro feeder zone; LRF., Le Roz feeder zone. Others

1082 abbreviations as in Fig. 2a.

1083

1084
48°00′N

<figure><img src="image_3.png" /><figcaption>(a)</figcaption></figure>

<figure><img src="image_4.png" /><figcaption>(c)</figcaption></figure>

<figure><img src="image_5.png" /><figcaption>(b)</figcaption></figure>

<figure><img src="image_6.png" /></figure>

50 km

<figure><img src="image_7.png" /><figcaption>(e)</figcaption></figure>
48°20′N

4°30′W

4°20′W

4°10′W

<figure><img src="image_8.png" /><figcaption>(a)</figcaption></figure>

<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /><figcaption>(b)</figcaption></figure>

<figure><img src="image_11.png" /></figure>

<figure><img src="image_12.png" /></figure>

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

<figure><img src="image_14.png" /></figure>

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

(Inferred) decollement
Thrust fault
Strike-slip fault
shear zone

(c)
<figure><img src="image_17.png" /></figure>

<figure><img src="image_18.png" /></figure>

<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /></figure>

<figure><img src="image_21.png" /></figure>

<figure><img src="image_22.png" /></figure>

<figure><img src="image_23.png" /></figure>
<figure><img src="image_24.png" /></figure>

<figure><img src="image_25.png" /></figure>

<figure><img src="image_26.png" /></figure>
<figure><img src="image_27.png" /></figure>

<figure><img src="image_28.png" /></figure>

<figure><img src="image_29.png" /></figure>
<figure><img src="image_30.png" /></figure>

<figure><img src="image_31.png" /></figure>

<figure><img src="image_32.png" /></figure>

<figure><img src="image_33.png" /></figure>

<figure><img src="image_34.png" /></figure>
<figure><img src="image_35.png" /></figure>

<figure><img src="image_36.png" /></figure>

<figure><img src="image_37.png" /></figure>
NW

<figure><img src="image_38.png" /></figure>

a. HIS cutting through S0

f. MIS parallel with S0

k. HIS oblique to S1

b. MIS intruded into a fold hinge zone

g. HIS parallel with S1

l. HIS/SIS transition

c. HIS/SIS transition

h. SIS intruded along a D3 shear thrust

m. SIS thrusted over and post-dating a HIS paralel with S1 (n)

d. SIS intruded along a D3 shear thrust

i. Highly strained HIS in a feeder zone

e. HIS parallel with S1

j. SIS parallel with S1

(a)

<table><thead><tr><th rowspan="2">Main tectonic events<br/>[≈ 320 – 310 Ma]</th><th colspan="2">Temporarily bracketed<br/>intrusive sheets [x 82]</th><th rowspan="2">Unconstrained<br/>sheets [x 19]</th></tr><tr><th>Fully bracketed<br/>[x 38]</th><th>Partly bracketed<br/>[x 44]</th></tr></thead><tbody><tr><td>Shearing<br/>D3<br/>Thrust<br/>Dextral<br/>strike-slip</td><td>Caro<br/>Larmor<br/>(x 7)</td><td></td><td></td></tr><tr><td>Folding<br/>D2</td><td>Squiffiec<br/>Le Roz<br/>(x 8)</td><td>(Missing<br/>upper limit)<br/>(x 18)</td><td></td></tr><tr><td>Syncleavage<br/>folding<br/>D1</td><td>South Château<br/>(x 5)</td><td>(Missing<br/>lower limit)<br/>(x 26)</td><td></td></tr></tbody></table>
<figure><img src="image_39.png" /></figure>

<figure><img src="image_40.png" /></figure>
<figure><img src="image_41.png" /></figure>

<figure><img src="image_42.png" /></figure>

<figure><img src="image_43.png" /></figure>

<figure><img src="image_44.png" /></figure>
<figure><img src="image_45.png" /></figure>
<figure><img src="image_46.png" /></figure>

<figure><img src="image_47.png" /></figure>

<figure><img src="image_48.png" /></figure>
<figure><img src="image_49.png" /></figure>

<figure><img src="image_50.png" /></figure>
<figure><img src="image_51.png" /></figure>

n = 94
<figure><img src="image_52.png" /></figure>

<figure><img src="image_53.png" /></figure>

<figure><img src="image_54.png" /></figure>

<figure><img src="image_55.png" /></figure>

<figure><img src="image_56.png" /></figure>

<figure><img src="image_57.png" /></figure>