Extensional flow produces recumbent folds in syn-orogenic granitoids (Padrón migmatitic dome, NW Iberian Massif)

Accepted Manuscript Extensional flow produces recumbent folds in syn-orogenic granitoids (Padrón migmatitic dome, NW Iberian Massif)

Rubén Díez Fernández, Luis Miguel Martín Parra, Francisco J. Rubio Pascual PII: DOI: Reference:

S0040-1951(17)30108-7 doi: 10.1016/j.tecto.2017.03.010 TECTO 127427

To appear in:

Tectonophysics

Received date: Revised date: Accepted date:

30 August 2016 17 February 2017 16 March 2017

Please cite this article as: Rubén Díez Fernández, Luis Miguel Martín Parra, Francisco J. Rubio Pascual , Extensional flow produces recumbent folds in syn-orogenic granitoids (Padrón migmatitic dome, NW Iberian Massif). The address for the corresponding author was captured as affiliation for all authors. Please check if appropriate. Tecto(2017), doi: 10.1016/j.tecto.2017.03.010

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ACCEPTED MANUSCRIPT Extensional flow produces recumbent folds in synorogenic granitoids (Padrón migmatitic dome, NW Iberian

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Massif)

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Rubén Díez Fernández1, Luis Miguel Martín Parra2, Francisco J. Rubio Pascual2

Departamento de Geodinámica, Universidad Complutense de Madrid. 28040 Madrid,

Instituto Geológico y Minero de España, 28760 Tres Cantos, Madrid, Spain

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Spain

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* Corresponding author at: Departamento de Geodinámica. Facultad de Geología, Universidad Complutense de Madrid, C/ José Antonio Novais, nº 2, 28040 Madrid,

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Spain. Tel.: +34 913944904; fax: +34 915442535

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E-mail address: [email protected] (R. Díez Fernández, corresponding author)

Keywords: Deformed granites; Recumbent folds; Extensional tectonics; Orogenic collapse; Migmatitic dome; Variscan orogen.

Abstract This contribution provides a case example on the generation of large-scale recumbent folds in syn-orogenic granitoids. We analyze the progressive reworking of extension-

ACCEPTED MANUSCRIPT related structures into later ones after a period of crustal thickening. The Padrón migmatitic dome formed after the climax of the Gondwana-Laurussia collision in the late Paleozoic. Petrostructural analysis carried out in the eastern flank of this dome reveals that extensional flow resulted in progressive exhumation of mainland Gondwana, which rested under peri-gondwanan allochthonous terranes and a suture

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zone during maximum crustal thickening. Exhumation proceeded up to upper crust

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levels (andalusite stability field) along with partial melting of the middle-lower crust

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and with the generation of granitoid laccoliths during an early extensional stage. Newlyformed lithological and mechanical anisotropies, such as the presence of variably-sized

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sheet-shaped bodies of syn-orogenic granitoids, provided a favorable rheological setting for fold nucleation during the intermediate stages of extension. In extending orogenic

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crust, whether recumbent folds occur after significant melt production depends on the lateral/vertical flow ratio, and on the orientation of deforming bodies with regard to

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kinematic/strain axes. We suggest that subhorizontal extensional flow dominated over

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vertical flow during the early and intermediate stages of the evolution of the Padrón dome. A component of vertical (diapiric) flow caused progressive tilting of the sheet-

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like bodies and obliquity respect to strain axes. This resulted in the development of regional-scale folds at the expense of syn-orogenic granitoids, such as in the case of the

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Portomouro recumbent synform. Extensional ductile flow was oblique to the trend of the orogen during the whole process, and directed to the NNW during the formation of recumbent folds. Non-coaxial shearing favored an (NNW-SSE) elongate shape for the syn-kinematic granitic massifs as well as the subsequent nucleation of recumbent folds. Deformation concentrated along discrete detachments during the late stages of extension.

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1. Introduction The gravitational and thermal reequilibration of thickened continental crust is associated with pervasive subhorizontal ductile shearing and the development of lowangle normal faults (extensional detachments) (e.g., Wernicke, 1981; Lister et al., 1984;

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Dewey, 1988; Malavieille, 1993; Brun and Van Den Driessche, 1994; Rey et al., 2001).

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In this process, lateral extensional flow increases metamorphic gradients (telescoping of

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crustal isotherms), and favors progressive overprinting by low-pressure/hightemperature metamorphism as well as partial melting of fertile layers, eventually giving

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way to the formation of migmatitic domes during stages dominated by vertical flow (e.g., Whitney et al., 2004b; Tirel et al., 2008; Vanderhaeghe, 2009; Kruckenberg et al.,

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2011). Ongoing deformation in this context has been recognized as an important factor during melt-segregation, melt-extraction, and melt-transport and intrusion in orogenic

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crust (Brown, 1994, 2013; Petford et al., 2000). The whole process, from melt-

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formation to intrusion, results in bodies of magma with a geometry conditioned on the stress field, i.e. sheet-like (laccoliths), oval-like, or wedge-shaped (e.g., Vigneresse,

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1995; Vigneresse et al., 1999; Aranguren et al., 2003; Talbot et al., 2004). Folds are one of the most recognizable geological structures. Their analysis

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provides rheological information about the rocks being folded as well as constraints on the history of deformation of orogenic belts (e.g., Hudleston and Treagus, 2010). Fold nucleation by buckling and subsequent amplification by superimposed flattening (usually under general shear) can explain the geometry of most folds (e.g., Hudleston, 1973; Sanderson, 1982; Ramsay et al., 1983; Ez, 2000; Bastida et al., 2014), although passive shearing is also considered as a possible mechanism for fold initiation (e.g., Ramsay, 1962). Fold nucleation is favored by the existence of mechanical instabilities

ACCEPTED MANUSCRIPT (e.g., rock bodies with unlike shape and irregularities, and with unlike effective viscosities), which would be amplified by layer-parallel or layer-oblique shortening to produce buckle folds (Ramberg, 1963; Treagus, 1973). Recumbent folds (Turner and Weiss, 1963; Fleuty, 1964; Ramsay, 1967) may be formed in a broad variety of tectonic settings and are usually found at some stage of

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tectonometamorphic evolution in metamorphic terranes. They may be localized in shear

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zones (e.g., Carreras et al., 2005), and/or widespread and defining the regional internal

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structure of a given terrane (e.g., Hatcher, 1981; Potts, 1983; Azor et al., 1994; Aerden and Malavieille, 1999; Díez Fernández et al., 2011). A combination of shortening and

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simple shear can explain the development of regional‐scale recumbent folds (Sanderson, 1982; Ramsay et al., 1983; Dietrich and Casey, 1989; Ez, 2000; Martínez Catalán et al.,

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2003). To explain the formation of recumbent folds in a compressional setting such as the Helvetic nappes, Ramsay (1981) proposed shortening along the hanging wall of an

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irregular thrust cutting across competent layers (see also Ramsay et al., 1983). However,

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it has been suggested that various amounts of overthrust displacement and superimposition of pure shear components on heterogeneous simple shear deformation

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are required to explain those nappes (Dietrich and Casey, 1989). In more internal zones of an orogen, differential flow can form passive folds. Then, contractional shearing

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concentrated along the reverse limb of an antiform would produce a ductile thrust that may amplify previous folds (standard model of nappe folds; Hatcher and Hooper, 1992). Recumbent folds can also be developed within sections of orogenic crust subjected to extension (Platt, 1982; Froitzheim, 1992; Orozco et al., 1997; Harris et al., 2002; Arango et al., 2013), and are rather common in migmatised areas, both at the micro- and meso-scale. Such folds are typically defined by the migmatitic banding, or

ACCEPTED MANUSCRIPT by meter-scale bodies of crystallized melt enclosed within. What is less common are cases of kilometer-scale recumbent folds formed coevally with the regional foliation that is found within migmatitic domes, so the role and information that can be extracted from this type of folds is yet limited. This contribution presents a field-based analysis of the structure of a granitoid

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massif and its host located in a migmatised area of the Variscan orogen, the Padrón

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dome (NW Iberian Massif). Here we provide an example of the complex folded

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structure developed in a kilometer-scale igneous massif that was formed during the thermal reequilibration of a collisional orogen. We discuss on the role of acting tectonic

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and/or gravity forces in the formation of folds.

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2. Geological setting

The Variscan orogen resulted from the collision of Gondwana and Laurussia in

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late Paleozoic times (Matte, 1991; Franke, 2000; Simancas et al., 2005; Martínez

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Catalán et al., 2007; Ribeiro et al., 2007; Ballèvre et al., 2009; Kroner and Romer, 2013; Díez Fernández et al., 2016). Fragments of this orogen are dispersed in southern and

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central Europe and occur as a series of crystalline massifs (Fig. 1). Among them, the Iberian Massif preserves a rather complete section of the orogen. In NW Iberia, due to

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the heterogeneous nature of deformation and metamorphism, the processes responsible for both the building (crustal thickening) and subsequent destruction (crustal thinning) of orogenic crust can be studied (e.g., Martínez Catalán et al., 2009). In NW Iberia (Fig. 2), tectonic events related to the early stages of the Variscan collision are evidenced in a set of allochthonous terranes, usually referred to as the allochthonous complexes (Arenas et al., 1986, 2016; Ribeiro et al., 1990; Martínez Catalán et al., 1997). These include two high-P metamorphic events affecting different

ACCEPTED MANUSCRIPT sections of Gondwanan continental crust, first in the Early Devonian (Ordóñez Casado et al., 2001; Fernández-Suárez et al., 2007) and then in the Late Devonian (Rodríguez et al., 2003; Abati et al., 2010). Continental subductions alternated with a transient stage of lithosphere extension and ocean-floor formation during the Early-Middle Devonian (Díaz García et al., 1999; Sánchez Martínez et al., 2011; Arenas et al., 2014b), thus

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contributing to the multi-stage character of the Gondwana-Laurussia collision (Arenas

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et al., 2014a).

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The translation of the allochthonous terranes on top of mainland Gondwana represents another major episode in the evolution of the orogen, being the main crustal

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thickening event recorded in the region. The transference occurred via thrusting (Ries and Shackleton, 1971; Ribeiro et al., 1990; Martínez Catalán et al., 1996) in the

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Carboniferous (Dallmeyer et al., 1997), and was associated with folding, both within the relative autochthon located underneath (Marquínez García, 1984; Farias et al., 1987;

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Ribeiro et al., 1990; Marcos and Farias, 1999; Díez Montes, 2007; Dias da Silva, 2014),

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and also within the allochthonous complexes located on top (Martínez Catalán et al., 1996; Díez Fernández and Martínez Catalán, 2009; Díez Fernández et al., 2011). The

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emplacement of the allochthonous terranes was coeval with the development of synorogenic sedimentary basins along the advancing front of tectonic nappes (e.g.,

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González Clavijo and Martínez Catalán, 2002; Dias da Silva et al., 2015; Martínez Catalán et al., 2016). These first deformation events propagated towards mainland Gondwana in the course of its collision with Laurussia (Dallmeyer et al., 1997). In the relative autochthon, early deformation was accommodated by upright to east-verging folds as well as by east- to southeast-directed thrusts (Ribeiro et al., 1990; Macaya et al., 1991; Díez Balda et al., 1995; González Clavijo and Martínez Catalán, 2002; Díez Montes, 2007; Díez Fernández et al., 2013; Rubio Pascual et al., 2013). The

ACCEPTED MANUSCRIPT development of all those structures was accompanied by Barrovian-type metamorphism, which reached the kyanite and sillimanite stability fields during this stage (0.5–1.0 GPa; Escuder Viruete et al., 1997; Díez Montes, 2007; Pereira, 2014; Rubio Pascual et al., 2015, 2016), and even eclogite facies conditions in some sections (~1.4 GPa; Barbero and Villaseca, 2000).

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Following the climax of tectonic thickening, a period of thermal relaxation and

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gravitational disequilibrium brought the orogenic crust to melting and extension

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(Martínez Catalán et al., 2014; Alcock et al., 2015). This stage is represented by voluminous magmatism and the development of migmatite-cored domes (e.g., Tormes,

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Martinamor, Gredos, Somosierra; Escuder Viruete et al., 1994; Díez Balda et al., 1995; Díaz-Alvarado et al., 2012; Arango et al., 2013; Rubio Pascual et al., 2013). These

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domes show pervasive overprinting of previous tectonothermal record by low-dipping tectonic foliation and conspicuous high-T and low-P metamorphism (e.g., Díez Balda et

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al., 1995; Escuder Viruete, 1998; Rubio Pascual, 2013). This latter phase shaped the

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hinterland of the orogen into a series of domes and structural basins. The allochthonous complexes now occupy the structural basins, while the relative autochthon is exposed in

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the cores of domes and antiforms (Fig. 2). After progressive lithosphere attenuation, the overcoming of gravity-driven

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extension (sourced from a gravitationally unstable orogen) by compressional forces (sourced from Gondwana-Laurussia convergence) resulted in a shift from crustal attenuation to a new phase of contraction and thickening (Díez Fernández and Pereira, 2016). This latter phase amplified the regional domes nucleated during previous deformation, and developed additional upright folds as well as strike-slip shear zones (Iglesias Ponce de Leon and Choukroune, 1980; Jiménez Ontiveros and Hernández Enrile, 1983; González Clavijo et al., 1991; Díez Fernández and Martínez Catalán,

ACCEPTED MANUSCRIPT 2012; Díez Fernández et al., 2013) dated at ca. 315–305 Ma (Capdevila and Vialette, 1970; Rodríguez et al., 2003; Valle Aguado et al., 2005; Gutiérrez-Alonso et al., 2015; Díez Fernández and Pereira, submitted). This new stage of compression is also represented by the folding of the orogenic belt about vertical axes, thus giving way to

et al., 2000, 2013; Aerden, 2004; Martínez Catalán, 2011).

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the development of plate-scale oroclinal bends in Iberia (Pérez-Estaún et al., 1988; Weil

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The Padrón migmatitic dome is a syn-convergent extensional system developed

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in the axial zone of the Variscan orogen that is exposed in the NW section of the Iberian Massif (Díez Fernández et al., 2012b). This dome is located between the allochthonous

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complexes of Órdenes and Malpica-Tui (Fig. 2). The northern part of the dome is flanked by a low-angle extensional detachment, the Pico Sacro detachment, which

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separates the overlying allochthonous units (suprastructure) from the underlying relative autochthon (infrastructure) (Martínez Catalán et al., 2002; Gómez Barreiro et al., 2010).

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The tectonometamorphic evolution of the different units that constitute the

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allochthonous ensemble is complex (e.g., Arenas et al., 2016). Specific data are available in numerous works about the basal allochthonous units (e.g., Martínez Catalán

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et al., 1996; Llana-Fúnez and Marcos, 2002; Rodríguez et al., 2003; Gómez Barreiro et al., 2010; Díez Fernández et al., 2011; López-Carmona et al., 2014), the allochthonous

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ophiolitic units (e.g., Díaz García et al., 1999; Arenas et al., 2007a, 2007b), and about the upper allochthonous units (e.g., Abati et al., 1999; Ábalos et al., 2003; Gómez Barreiro et al., 2007). The emplacement of the allochthonous units towards the east and onto their relative autochthon, which now crops out in the dome infrastructure, has been dated at ~340 Ma (Martínez Catalán et al., 1996; Dallmeyer et al., 1997). Deformation (thickening) within the autochthon via folding (e.g., Macaya et al., 1991; Díez Balda et al., 1995; Díez Fernández et al., 2013) started earlier (ca. 354-347 Ma; Bea et al., 2009;

ACCEPTED MANUSCRIPT Rubio Pascual et al., 2013). Much of the penetrative deformation of the allochthon occurred before emplacement, so thrusting juxtaposed pieces of orogenic crust bearing their own previous tectonic record (Díez Fernández et al., 2016 and references therein). The metamorphic evolution of the relative autochthon has been poorly constrained in previous works. It includes an initial stage of deformation accompanied

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by amphibolite facies metamorphism followed by an exhumation path up to the upper

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crust through partial melting conditions for metapelites (e.g., Díez Fernández, 2011;

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Díez Fernández et al., 2012b). Nearby exposures of the relative autochthon to the west of the Padrón dome suggest a similar evolution (Gil Ibarguchi, 1982).

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The main foliation that is observed both in the dome suprastructure (allochthonous complexes) and in the dome infrastructure (relative autochthon),

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contours the domal structure of Padrón (Díez Fernández, 2011). Some of the granitoids that occur within the dome infrastructure cut across the main foliation that defines the

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dome and represent late-kinematic intrusions. The primary contacts of some other

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granitoids are subconcordant with the main foliation, even in those zones characterized by low-dipping foliation. This work focuses on the structure and tectonic evolution of a

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granitic massif and its metasedimentary host rocks located in the eastern limb of the

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Padrón dome, west of Santiago de Compostela city (Fig. 2).

3. Lithostratigraphy, petrography and relationships between rock types The study area can be divided in an allochthonous and a relative autochthonous domain (Figs. 3a and 3b). The allochthonous units include ortho- and paragneisses, schists, eclogites, and amphibolites, while the relative autochthon consists of metasedimentary rocks and various types of granitoids. The metasedimentary rock sequence of the relative autochthon includes mica schists, quartz-rich schists,

ACCEPTED MANUSCRIPT paragneisses, migmatitic metasedimentary rocks, and minor quartzites and graphitebearing schists. It corresponds to the so-called Paraño Group (Marquínez García, 1981), which is a member of the Parautochthon of NW Iberia (Farias et al., 1987; Barrera et al., 1989). Based on its detrital zircon input, a West Africa Craton provenance and a Gondwanan affinity have been proposed for this sequence (Díez Fernández et al.,

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2012a).

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The metasedimentary rocks of the relative autochthon show an increase in

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metamorphic grade. Some of those rocks have been partially to almost fully melted (metatexites and diatexites; following Brown, 1973), and there is a progressive increase

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in the abundance of leucosomes, granitic dikes and sills from the boundary with the allochthonous units to the inner parts of the relative autochthon. The rocks of the

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allochthonous domain do not show evidence of equivalent melting processes. Bedding is sometimes preserved in the relative autochthon, although a tectonic and/or a

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migmatitic banding is usually the only planar fabric identifiable.

3.1 Metasedimentary rocks

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The mineral assemblage in non-migmatitic metasedimentary rocks of the relative autochthon includes quartz, muscovite, biotite, plagioclase, and minor amounts of

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sillimanite, andalusite, garnet, tourmaline, zircon, apatite and opaque minerals. The shape-preferred orientation of planar and elongated minerals in these rocks, together with that of quartzo-feldspathic domains, define a penetrative foliation (Fig. 4a). Two main petrological groups can be distinguished in the migmatitic metasedimentary rocks, the metatexites and the diatexites. The mineral assemblage in the metatexites consists of quartz, biotite, plagioclase, K-feldspar, sillimanite, andalusite, and accessory amounts of garnet, cordierite, tourmaline, apatite, zircon, and

ACCEPTED MANUSCRIPT opaque minerals (Fig. 4b). Muscovite may be present as flakes grown over the migmatization structures. Chlorite after biotite and muscovite after feldspar or aluminum silicate are common low-temperature transformations, while fibrolitic sillimanite can be enclosed within muscovite and anhedral andalusite (Fig. 4c), or be crosscut by subeuhedral crystals of andalusite (Fig. 4b and 4d).

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Metatexites are made of three main components. The leucocratic component is

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defined by thin layers, branching veins and minor pods with medium- to coarse-grained

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quartz, feldspar, plagioclase and minor (peritectic?) garnet and biotite (leucosome; Fig. 4e). The melanocratic components alternate layers enriched in biotite, and sometimes in

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sillimanite (melanosome), with layers richer in medium- to fine-grained quartz, feldspar and plagioclase (paleosome). Sillimanite and andalusite may occur in the melanosomes,

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which have a well-defined foliation marked by planar and elongated minerals, such as biotite and sillimanite (Fig. 4b), as well as by the shape-preferred orientation of quartz,

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plagioclase and feldspar. Most of the boundaries of the leucosomes are parallel to the

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main foliation and/or bedding (Fig. 4e), but some crosscutting relationships are also observed. In fact, quartz, plagioclase, K-feldspar and biotite in the leucosomes may

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show shape-preferred orientation parallel to the main foliation and usually display evidence of solid-state deformation, including internal deformation and

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recrystallization, grain-size reduction, foliation anastomosing, microcline twinning, myrmekite, boudinage of competent minerals, and development of discrete shear zones within the leucosome (e.g., Fig. 4f). Diatexites of the relative autochthon contain enclaves of metatexite and consist of rather variable amounts of quartz, plagioclase, K-feldspar, biotite, and minor amounts of tourmaline, apatite, rutile, zircon, garnet, and opaque minerals. These diatexites show a primary granitic texture, but quartz, biotite, euhedral/subeuhedral feldspar and

ACCEPTED MANUSCRIPT plagioclase are deformed at solid-state conditions and usually show preferred orientation, thus defining a foliation. The orientation of some quartz and plagioclase is random and their shape tends to be equant.

3.2 Granitoids

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The granitic massifs of the study area (Fig. 3a), as well as other granitic bodies

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of the Padrón dome with similar characteristics are Variscan in age (320-300 Ma;

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Bellido et al., 1992; Rodríguez et al., 2003, 2007). They are affected by Variscan deformation to some extent (see below), and so they are considered as syn-orogenic.

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They essentially occur in the autochthonous domain and can be divided into synkinematic and late-kinematic granites (Capdevila et al., 1973; Corretgé, 1983; Castro et

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al., 2002).

Syn-kinematic granites include two main types of rocks: a dominant granite-

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diatexite ensemble, and patches of medium- to coarse-grained two-mica granites. Both

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of them contain quartz, K-feldspar, plagioclase, biotite, muscovite, and minor amounts of apatite, zircon, rutile, sillimanite and opaque minerals. The textures in the granite-

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diatexite ensemble are very variable and range from most common allotriomorphicgranular to hypidiomorphic (even porphyritic for feldspar), with grain sizes spanning

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from fine to coarse. Mineral proportions in the latter vary significantly from one outcrop to the next, and patches or elongate bodies of metatexite and diatexite are common in some areas. For all that, the granite-diatexite ensemble has been classically referred to as “inhomogeneous granites” in the literature. The medium- to coarse-grained two-mica granites are much more homogeneous, both in grain-size (typically medium-grained) and in the relative proportion of minerals. Yet, the range of textures is similar to that of

ACCEPTED MANUSCRIPT the granite-diatexite ensemble (allotriomorphic-granular to hypidiomorphic), including porphyritic varieties. Enclaves of metatexite are generally absent. The syn-kinematic granites exhibit facies with preferred orientation of minerals and facies where orientation is either too vague to be recognized with the naked eye or lacking. When distinguishable, the main planar fabric that can observed is consistently

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parallel to that of the adjacent diatexites, metatexites, and schists (Fig. 3a), but never too

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intense as to produce incipient compositional banding. At the mesoscale, biotite,

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muscovite, and elongate feldspar define the fabric (Fig. 5a). At the microscale, subgrain boundaries of quartz, plagioclase and feldspar crystals, together with recrystallized

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aggregates of muscovite and/or biotite located around the boundaries of feldspar and plagioclase, add to the textural elements defining the main foliation (Fig. 5b). The latter

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examples suggest solid-state deformation, although the presence of aligned euhedral crystals of feldspar that are not internally deformed (or just weakly deformed) and that

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are imbricated with other similar grains, suggest that the planar fabric is probably a

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combination of magmatic flow and superimposed solid-state flow taking place as the magma reached its solidus (Paterson et al., 1989; Vernon, 2000). Within the syn-

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kinematic group, the foliation seems more penetrative and widespread for the case of the granite-diatexite ensemble.

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The late-kinematic group is represented by fine- to medium-grained two-mica granites. These have a similar mineral composition to the granitoids described before, but never show penetrative foliation. However, a scattered and moderate alignment of elongate euhedral feldspar and biotite crystals within a non-oriented matrix of quartz, plagioclase and feldspar is the only planar fabric that may be seen. The bodies of latekinematic granites cut the main foliation of the study area (RP-1; Fig. 3a), as well as some of the straight contacts of the syn-kinematic granites (RP-2; Fig. 3a) and the

ACCEPTED MANUSCRIPT boundary between the allochthon and the relative autochthon (RP-3, 4; Fig. 3a). The latter boundary cuts the massif of syn-kinematic granites (RP-5; Fig. 3a), providing additional evidence of the older age of the syn-kinematic granites relative to the latekinematic ones. The contacts between the late-kinematic granites and the rest of the sequence are

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sharp, whereas the contacts of the granite-diatexite ensemble are usually transitional

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into the medium- to coarse-grained two-mica granites, and more importantly, into the

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hosting migmatitic metasedimentary rocks. In these metatexites, there is a general increase in the proportion of lenses of leucogranite (e.g., Fig. 5c; dated at 317 ± 3 Ma by

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Rodríguez et al., 2003), and in the percentage of leucosome towards the main body of granite. Near the anatectic granite, pre-migmatization structures are frequently not

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preserved and the metatexite migmatites turn progressively into diatexite migmatites. Away from the main anatectic granite massif, partially melted rocks also exist, but they

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do not seem to define continuous layers and occur irregularly distributed all through the

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study area. For them, the percentage of leucosome is usually smaller than in the migmatites located next to anatectic granites.

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Given the diffuse and transitional nature of the contacts between the various types of migmatites and the rest of the rocks, the cartographic boundaries of the granite-

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diatexite ensemble presented in Figure 3a represent virtual surfaces where the transition from migmatite to anatectic granite occur. In some cases, these contacts separate diatexite migmatite from metatexite migmatite, although there exist other cases where such a contact is clearly not as sharp as it could be deduced from the map, even at the regional scale, since some patches of migmatite metatexite are still observable well inside the main body of anatectic granite. On the other hand, the anatectic granite may also be directly in contact with either metatexites or even non-migmatitic rocks.

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4. Deformation history Three main phases of deformation affected the relative autochthon of the study area (D1, D2, and D3). Due to strong overprinting, and the prograde nature of subsequent metamorphism, the record related to D1 (S1) is very incomplete. D2 produced the main

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foliation in the relative autochthon (S2), whereas D3 is responsible for the folded

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geometry of S2 at the regional scale and the development of local crenulation cleavage

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(S3). All these structures are cut by subvertical faults (with small offsets), which can be grouped in two main sets, one trending NW-SE and another one trending NE-SW (Fig.

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3a).

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4.1 D1 (S1)

The preservation of a foliation (S1) associated with a first phase of deformation

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(D1) in the relative autochthon of the study area is diffuse due to intense subsequent

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deformation (D2). The early-D2 fabrics preserved in the metapelites of the relative autochthon are represented by a tectonic banding with quartz-rich and mica-rich layers

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(Fig. 5d). This kind of banding is usually produced by metamorphic differentiation (solution + recrystallization + mechanical differentiation) during progressive

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deformation (e.g., Eskola, 1932; Ghaly, 1969), and/or typically by differentiation along cleavage planes formed in a previously foliated rock (e.g., Talbot and Hobbs, 1968; Gray, 1977), i.e. a rock containing a S1 fabric. S1 is better preserved in the northern part of the Padrón dome (located out of the study area). There, S1 is preserved within D2-porphyroblasts (andalusite) grown in the metapelitic rocks of the relative autochthon. S1 includes quartz, biotite, white mica, garnet and staurolite, suggesting mid-pressure amphibolite facies metamorphism (Díez

ACCEPTED MANUSCRIPT Fernández, 2011; Díez Fernández et al., 2012b). No major D1 structures are observed within the relative autochthon of the study area due to superimposed deformation. In less structurally-reworked sections of the relative autochthon located more to the east, D1 is related to east-verging folds and thrusts (Marquínez García, 1981; Aller y Bastida,

PT

1996; Marcos y Farias, 1999; Díez Montes, 2007; Rubio Pascual et al., 2013).

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4.2 D2 (S2)

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The main foliation of the metasedimentary rocks of the study area (S2) ranges from a schistosity to a gneissic banding, and both types may coexist with subparallel

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local migmatitic banding. The latter can be also widespread and define the entire planar structure in certain sectors. S2 is a composite planar fabric (following Tobisch and

MA

Paterson, 1988) that derives from two foliation-forming events. The two planar fabrics that make S2 can be observed either subparallel or

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oblique to each other. Recognition of these two components of S2 is easier when

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leucosomes are present. When the two foliations are parallel, the leucosomes and S2 are concordant (e.g., Fig. 4e). In these cases, evidence of previous mineral fabrics and the

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early-D2 nature of the leucosomes is best preserved at the microscale. In the one hand, the schists and paragneisses (even the paleosome) display a tectonic banding, alternating

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mica-rich (biotite and white mica) and quartz-rich layers, which can be transected by a penetrative cleavage (S2; Fig. 5d). Such tectonic banding would correspond to vestiges of an early-D2 foliation, most likely a S1 + proto-S2 foliation. On the other hand, the leucosomes are deformed under solid-state conditions (Fig. 4f), so the generation of melt they account for must have been followed by deformation after crystallization. When the two D2-fabrics are oblique, S2 appears as a crenulation cleavage. Relicts of bedding, migmatitic banding (Fig. 5e), and high-T mineral fabrics (Fig. 4d) are

ACCEPTED MANUSCRIPT similarly affected by such cleavage, thus suggesting that partial melting, at least some of it, is essentially previous to the final development of S2. In this regard, the planar primary nature of the leucosomes in the study area could be also taken as indicative of a pre-S2 planar anisotropy, such as a S1 + proto-S2 foliation, which would have conditioned local melt segregation, migration, and accumulation (e.g., Sawyer, 2001).

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The main mineral assemblage that can be inferred for the metasedimentary rocks

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during the early-D2 stage includes quartz, plagioclase, K-feldspar, biotite, sillimanite,

SC

garnet, tourmaline and opaque minerals. The mineral assemblage associated with the later development of S2 consists of quartz, plagioclase, biotite, muscovite, andalusite,

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cordierite, and opaque minerals.

The main foliation in the syn-kinematic granitic rocks contains quartz, feldspar,

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plagioclase and mica that are preferentially oriented, defining a single foliation (S2) throughout the granite-diatexite ensemble (Fig. 5a). Those minerals can be aligned too,

D

and may define a mineral and stretching lineation together with elongated enclaves and

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restites. This lineation is parallel to that observed in the metasedimentary host nearby. A comparable planar and linear fabric is absent in the late-kinematic granites.

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S2 is not homogeneous regarding associated structures, penetrativeness, and intensity of associated non-coaxial shearing. Besides folds and crenulation cleavages, S2

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may also appear as a discrete shear band cleavage (SC-C'). The angle between older foliation and the shear bands is consistently less than 45º, and deformation seems more intense (more grain-size reduction) compared to neighboring rocks. However, the thickness of these shear bands is typically less than 3 cm, their lateral extension is limited, and their distribution is random at the regional scale. S2 is generally low- to moderate-dipping, although it can also show subvertical dips. Changes in dip values and dip-direction of S2 are due to later deformation (see

ACCEPTED MANUSCRIPT below). S2 is axial planar to NE-SW trending recumbent folds (Fig. 3b), observable both at the micro- (Fig. 4d) and meso-scale (Fig. 5e). S2 minerals, especially mica, are aligned to define a mineral lineation (Fig. 5f). Its trend coincides with the stretching of quartz and feldspar grains, as well as with the long axis of small lenses of leucosome, so it is considered a stretching lineation (L2s). L2s trends NNW-SSE and is oblique to F2,

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i.e. at an angle less than 90° counterclockwise in relation to F2 (Fig. 3d). D2 kinematics

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is top-to-the-NNW, as inferred from the L2s trend combined with asymmetric S2

SC

structures such as σ-type objects (porphyroclasts and mineral aggregates), mica-fish, SC

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composite foliations, and C´ fabrics.

4.3 D3 (S3)

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S2 is affected by upright folds (D3) that trend NNW-SSE (Fig. 4e). Their fold axes (F3) plunge gently to the NNW, with local changes in the plunge direction (Fig.

D

3c). The strike of their axial planes (S3) range between NW-SE and NNE-SSW (mean

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orientation at NNW-SSE; Fig. 3c), and their dip-direction can be to the E or W. The amplitude of these folds is variable, up to several km, and they are generally open. D3

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folds may show an associated subvertical crenulation cleavage (S3) defined by quartz + chlorite + muscovite and the reorientation of S2 minerals (Fig. 4e). Some of the minerals

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defining S3 (especially muscovite) can be aligned with the axes of the upright folds. The planar fabric of the late-kinematic granites is usually high-dipping and oblique to the main planar fabric in the surrounding rock hosts (S2), including that of the granite-diatexite ensemble (Fig. 3a). The foliation in the late-kinematic granites is not penetrative, as opposed to that of the granite-diatexite ensemble. D3 upright folds do not affect the foliation of the late-kinematic granites. Instead, that foliation is parallel to the

ACCEPTED MANUSCRIPT axial planes of D3 folds and to S3 crenulation cleavage, so it has been tentatively considered as an equivalent planar fabric.

5. Major structures 5.1 Portomouro recumbent synform

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The penetrativeness and intensity of D2 in the syn-kinematic granites allow

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applying the methodology implemented by Díez Fernández and Martínez Catalán

SC

(2009) to infer the regional structure of deformed igneous massifs. S2 defines a major open upright D3 fold, the Milladoiro antiform, which is accompanied by minor upright

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folds (Fig. 3b).

The massif of syn-kinematic granites displays sinuous contacts with their

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metamorphic host, but the relationship between these contacts and the main foliation (S2) in the host and granites varies from layer-parallel to oblique or perpendicular.

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Crosscutting relationships are observed in zones where the map pattern of the contacts

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of the syn-kinematic granites resembles that of a hinge of a recumbent fold. Such relationship can be observed both in the case of a granite-metasedimentary rock contact

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(RP-6 to 10; Fig. 3a) and between the two types of syn-kinematic granites (RP-7, and 11 to 13; Fig. 3a). On the other hand, some of the straight and curved contacts of the

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granitic massif follow the strike of S2, and consequently represent zones of the granitic massif controlled by its D3 upright structure. In these cases, straight contacts would represent D3 fold limbs (e.g., RP-14; Fig. 3a), while curved contacts would be D3 fold hinge zones. The D3 Milladoiro antiform provides good examples of the latter case, both for its northern and southern terminations. In the two cases, the antiform plunges to the north (Figs. 3a and 3c), i.e. in the direction that the folded contact of the granitic massif closes (e.g., RP-15 and 16; Fig. 3a). Consequently, the northern contact of the syn-

ACCEPTED MANUSCRIPT kinematic granitic massif would represent its upper structural limit, whereas the southern contact would be the lower one. The rest of the curved contacts of the synkinematic granitic massif may be interpreted as either hinge zones of recumbent folds or terminations of the granitic massif. If we are dealing with regional recumbent folds, those terminations could represent the limbs.

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A geometrical correlation established between the curved contacts in both limbs

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of the D3 Milladoiro antiform reveals a (neutral) recumbent fold structure defined by the

SC

massif of syn-kinematic granites (Fig. 3b). The intrados would be located to the north, and represented by a hairpin-like shape of the contact between the granite-diatexite

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massif and the metasedimentary host (RP-6; Fig. 3a). This contact is bent, and so is the hinge zone of the recumbent fold, which exhibits hook-type interference with the

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Milladoiro antiform. The extrados would be located to the south, and symmetrically repeated over the two limbs of the Milladoiro antiform (RP-8-9; Fig. 3a). The upper

D

limb of the recumbent fold would be well exposed in the northern part (RP-16; Fig. 3a),

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where it defines the northern closure of the Milladoiro antiform and seems to end towards the west (RP-17; Fig. 3a). The upper part of the lower limb is partly exposed

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west of Portomouro (RP-18; Fig. 3a), whereas its lower part defines the southern closure of the Milladoiro antiform (RP-15; Fig. 3a).

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The recumbent fold structure can be deduced both for the granite-diatexite ensemble and, less clearly, for the syn-kinematic two-mica granites. In the latter case, the recumbent structure is more diffuse, especially for the northwestern occurrences of these rocks. Yet, a broad correlation of hinge zones can be proposed across the D3 upright structure of the Milladoiro antiform (RP-11 correlates with RP-13, and RP-12 correlates with RP-13; Fig. 3a). Moreover, the two main exposures of syn-kinematic two-mica granites can be correlated across the two limbs of the antiform, so the two

ACCEPTED MANUSCRIPT occurrences would account for a single granitic body. They lie into the hinge zone of the main recumbent fold. But in the southeastern part (RP-8; Fig. 3a), they occupy the extrados, while in the northwestern part they are located closer to the intrados (RP-11 and RP-12; Fig. 3a). This suggests not only obliquity between the axis of the major recumbent fold and the body of syn-kinematic two-mica granites, but also a NNW-SSE

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elongate shape for that igneous body. To this figure, the whole massif of syn-kinematic

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granitoids seems to be a lens-like igneous body that stretches for a minimum of 24 km

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in a NNW-SSE direction. This is constrained by its termination to the southeast in the hinge zone of the recumbent fold structure (RP-8-9; Fig. 3a), and by its termination to

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the northwest, at least in the upper limb of the recumbent fold (RP-17; Fig. 3a. Although the actual length of the massif in the E-W direction cannot be constrained (the

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massif pinches in on both limbs of the Milladoiro antiform), a minimum of 8 km can be assumed based on the E-W thickness of the massif in map view. However, the

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stretching lineation observed in the main foliation of the region trends NNW-SSE (L2s;

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Fig. 3d), so a lens-like geometry stretched in that direction is favored for the granitic massifs. Note that observable extent of the massif in the NNW-SSE direction is three

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times the one visible in the E-W direction. The existence of this fold is also supported by observations at outcrop-scale.

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Minor D2 recumbent folds (F2) affecting migmatites and granites can be observed in the region (Figs. 4d and 5e). The trend of this particular set of folds ranges from N-S to NESW, with a mean regional trend of N36°E (Fig. 3d). Such variation in the fold axis trend is similar to the dispersion that can be inferred from geometrical correlation of recumbent fold hinge zones across the two limbs of the Milladoiro antiform (see white dashed lines indicating the correlation in Fig. 3a). Local D2 minor fold asymmetries are compatible with the regional structure deduced from the map, although the number of

ACCEPTED MANUSCRIPT observations is relatively limited (18 in total, 11 with clear asymmetric folds). Since the regional foliation (S2) is axial planar to the D2 folds at the outcrop and thin-section scale, and a similar geometrical relationship is deduced between S2 and the major recumbent folds inferred from map analysis, recumbent folding is considered one of the expressions of D2 at the macro-scale. According to the top-to-the-NNW kinematics

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observed in S2 (see section 4.2; Fig. 3d), the main recumbent fold nucleated in the syn-

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kinematic granitoids, together with the rest of parasitic folds that formed along with it

SC

(e.g., RP-10; Fig. 3a), are considered as a synform (i.e. a member of a NNW-vergent

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structure), and will be referred to as Portomouro synform.

5.2 Pico Sacro detachment

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The Pico Sacro detachment is a regional-scale ductile/brittle shear zone that cuts the main foliation (S2) and the D2 recumbent folds developed in the relative autochthon

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(RP-5; Fig. 3a). This detachment is affected by the D3 Milladoiro antiform, which is

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responsible for its current subvertical dip (Gómez Barreiro et al., 2002, 2010; Martínez Catalán et al., 2002). Note that this detachment trends NNW-SSE in the northern part,

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and changes its trend to a N-S direction in the southern part, near the hinge zone of the Milladoiro antiform. Therefore, the Pico Sacro detachment can be framed into the last

AC

stages of D2 deformation.

The thrust pile before the Pico Sacro detachment was formed included the Autochthon and the Parautochthon in the lowermost part. On top of them were located, from bottom to top, the Basal, Ophiolitic and Upper Allochthonous units. The stacking of metamorphic units was carried out by east- to southeast-directed thrusts (Martínez Catalán et al., 1997, 2007), so a primary west-dipping geometry is expected for the thrust planes. To the east, the upper part of the footwall of the Pico Sacro detachment is

ACCEPTED MANUSCRIPT occupied by units located upper in the pre-Pico Sacro tectonic pile (Basal and Ophiolitic Allochthonous units), and to the west, the upper part of the footwall is defined by the lower structural units, i.e. the Parautochthon and Autochthon (Fig. 3). This geometry implies a west-dipping component for the Pico Sacro detachment. On the other hand, the western section of the Pico Sacro detachment defines the basal contact of the

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northern half of the Malpica-Tui Complex (Díez Fernández, 2011). This complex is

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preserved as a klippe in a late upright synform that folds the Pico Sacro detachment too.

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The regional plunge of the upright fold structure of the complex is to the north in the northern half (Díez Fernández et al., 2011), and to the south in the southern half (Díez

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Fernández and Martínez Catalán, 2009). The Pico Sacro detachment do not occur in the southern half of the Malpica-Tui Complex, what points to a north-dipping component

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for the Pico Sacro detachment (Díez Fernández et al., 2012b). Kinematic criteria associated with the movement of the Pico Sacro detachment

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indicate top-to-the-northwest sense of shear (Gómez Barreiro et al., 2002, 2010; Díez

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Fernández et al., 2012b). Since it was a detachment inclined to the north and to the west and moved its hanging wall to the northwest, the Pico Sacro can be interpreted as a low-

AC

D3 folding.

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angle normal fault, i.e. an extensional detachment (Martínez Catalán et al., 2002) before

6. Discussion on the tectonic evolution The petrostructural analysis performed in this work, together with the regional setting, suggest the following sequence of major tectonometamorphic events for the relative autochthon (Parautochthon) that crops out in the eastern flank of the Padrón migmatitic dome (Fig. 6): (i) tectonic juxtaposition of the allochthonous complexes on top of the relative autochthon causing increasing pressure on the latter (D1); (ii)

ACCEPTED MANUSCRIPT exhumation of the relative autochthon via tectonic denudation and erosion acting over an evolving migmatitic dome (D2); and (iii) limited crustal thickening and exhumation due to superimposed upright folding (D3). This general sequence is common to most of the migmatite domes of the Iberian Massif, both in NW Iberia (Martínez Catalán et al., 2002, 2003; Gómez Barreiro et al., 2010; Díez Fernández et al., 2012b), Central Iberia

PT

(Escuder Viruete et al., 1994, 1998; Barbero, 1995; Díez Balda et al., 1995; Díaz-

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Alvarado et al., 2012; Díez Fernández et al., 2013; Rubio Pascual et al., 2013; Díez

SC

Fernández and Pereira, 2016), and SW Iberia (Pereira et al., 2009).

It is expected that the juxtaposition of the allochthonous complexes onto the

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relative autochthon has favored metamorphic reactions and the formation of mineral fabrics in the latter. The early metamorphic evolution in the relative autochthon was

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related to the formation of mid-pressure mineral parageneses (M1). This part of the metamorphic path is not clearly preserved in the study area, but it can be deduced from

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the >15km thick pile of allochthonous units that must have rested on top by the climax

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of crustal thickening. Such value can be inferred from the minimum thickness of the whole overlying allochthonous thrust pile that is observed today around the Padrón

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dome (Fig. 2), and is consistent with thermobarometric estimations for the onset of D2 in close easterly areas of the Parautochthon (0.75 GPa; Rubio Pascual et al., 2015).

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Besides, in other equivalent sections of the uppermost relative autochthon located to the north, along the coastal section west of the Órdenes Complex, the metasedimentary rocks still preserve some relict assemblages with quartz + biotite + white mica + garnet ± staurolite (Díez Fernández, 2011; Díez Fernández et al., 2012b). Even kyanite-bearing assemblages are also present in close easterly areas of the Parautochthon. The development of the main foliation (S2) in the study area was progressive, and took place as the relative autochthon underwent decompression. This is typified (i)

ACCEPTED MANUSCRIPT by the deformation of an early-D2 migmatitic and tectonic banding during the development of S2 (Figs. 4f and 5d), (ii) by the growth of D2-andalusite after sillimanite in the metasedimentary rocks (Figs. 4b, 4c, and 4d), and (iii) because D2-andalusite grew after maximum crustal thickening (D1), which provided lithostatic load (or pressure) in the relative autochthon (at least) up to ~0.75 GPa (staurolite stability field).

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Melt-production curves for metapelites were cut essentially during decompression, as

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indicated by the presence of (limited) garnet within early-D2 leucosomes (Fig. 4f; Spear

SC

et al., 1999). Partial melting of the metasedimentary rocks occurred under T < 800°C, within the stability field of biotite (note the absence of pyroxene; Vielzeuf and

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Holloway, 1988), and largely at the expense of muscovite, plagioclase, and quartz (e.g., Thompson and Tracy, 1979), suggesting T > 600ºC. The growth of metamorphic S2-K-

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feldspar in the metasedimentary rocks suggest T > 650ºC. We cannot ruled out the possibility that some partial melting occurred during D1

D

crustal thickening. However, metamorphic conditions in sections that occupied a

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seemingly equivalent position to the study area during D1 reached the staurolite stability field for metapelites. Therefore, melts, if present, should have derived from deeper

CE

sources. So far, there is no data in NW Iberia supporting extensive partial melting older than 340 Ma, which is considered the reference age for the transition from D1

AC

(emplacement of the allochthonous complexes; Martínez Catalán et al., 1996; Dallmeyer et al., 1997) to D2 in the relative autochthon. According to the tectonomagmatic, metamorphic, and structural record, the exhumation of the relative autochthon occurred in three main steps during D2 (see below).

6.1. Early-D2 stage: onset of gravitational collapse

ACCEPTED MANUSCRIPT The early exhumation of the relative autochthon comprises the processes developed between the end of D1 (Fig. 6a) and the onset of D2 recumbent folds. This stage included subhorizontal flow, partial melting and emplacement of large sheet-like massifs of granitoid in the relative autochthon (Fig. 6b). The Portomouro synform was nucleated from an irregular, yet flattened and probably stretched igneous body, as

PT

constrained from the structural analysis presented in section 5. This idea is also inspired

RI

by numerous observations on the lensoidal geometry of small bodies of syn-kinematic

SC

granite at the outcrop scale (e.g., Fig. 5c), and supported by the remarkable parallelism between the main foliation (S2) and the boundaries of the largest massif in map view

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(Fig. 3a).

There are no isotopic age constrains for the granitic massifs affected by

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recumbent folding, and a pre-D2 (i.e. D1-related) age could explain a pre-fold, sheet-like structure in the granitoids, either acquired during D1 or early-D2. However, these

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granitoids do not show folded pervasive planar fabrics, and their main foliation (S2)

PT E

does not present evidence for a composite origin, as it would be expected for pre-D2 massifs. Note that early-D2 deformation produces a penetrative planar anisotropy in the

CE

region (proto-S2). That anisotropy is folded into the Portomouro synform, as observed in the metasedimentary rocks, and should be equally folded in the granitoids if they

AC

were pre-D2. The sheet-like nature of the granitic massif, as well as the lensoidal shape of minor patches of granite lying within the migmatised rocks that are now affected by D2 recumbent folds, must have been acquired at early D2. Such kind of structure is typical for syn-kinematic granitoids (Castro, 1987; Brown, 1994; Vigneresse, 1995; Brown and Solar, 1999; Vigneresse et al., 1999; Díaz-Alvarado et al., 2012), and suggests subhorizontal tectonic flow during their genesis and intrusion.

ACCEPTED MANUSCRIPT The tectonic fabrics formed during this phase were progressively reworked by later ductile deformation in the course of D2. Consequently, we have no reliable kinematic indicators that we could use to constrain a potential direction of movement for this particular stage. However, the NW-SE stretched shape that the granite-diatexite and two-mica granite massifs may have had before they were affected by the

PT

Portomouro synform (see discussion in section 5), would suggest that dominant tectonic

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flow at this initial stage was oblique to the trend of the orogen.

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The interpretation of D2 subhorizontal flow is different whether linked to reverse or to normal fault movement. Unfolding of S2 and related structures does not provide a

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clear primary dipping direction of that fabric before D3. Kinematic criteria associated with the development of S2 across the whole Padrón dome define a regional divergent

MA

flow during D2. The northern section of the Padrón dome flowed to the northwest, whereas the southern section moved to the southeast (Díez Fernández et al., 2012b). At

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a large scale, the regional foliation of the dome (S2) pitches towards the foreland, under

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the upper-crustal sections located east (Fig. 2). This feature is compatible with the geometry of a normal fault, and together with the divergent flow pattern governing D2

CE

constitute typical features of extensional settings. Following the emplacement of the allochthonous complexes on top of their

AC

relative autochthon (D1) at ~340 Ma (Martínez Catalán et al., 1996; Dallmeyer et al., 1997), extensional tectonics dominated in NW Iberia (D2). The age of the oldest extensional detachment related to the collapse of the overlying allochthonous thrust pile and the exhumation of the relative autochthon in the Padrón dome (Bembibre-Ceán detachment; Díez Fernández et al., 2012b) is 337 ± 3 Ma (López Carmona et al., 2014). The age of some syn-D2 leucogranites occurring in the relative autochthon is 317 Ma (40Ar/39Ar; Rodríguez et al., 2003). The major structures associated with S2 are cut by

ACCEPTED MANUSCRIPT the Pico Sacro extensional detachment, which is folded by late D3 upright folds dated at ca. 315-305 Ma (Capdevila and Vialette, 1970; Rodríguez et al., 2003). From 337 Ma to 315 Ma, there is no evidence of crustal thickening (e.g., metamorphic inversion or pressurization) in the relative autochthon of the Padrón dome. The coexistence of abundant granitoids and low-pressure metamorphism

PT

(andalusite-sillimanite) is common in extensional tectonic setting (e.g., Wickham, 1987;

RI

De Yoreo et al., 1991; Whitney and Dilek, 1998; Whitney et al., 2004a). It is also found

SC

during the extrusion of metamorphic nappes bounded by a top normal fault and a lower thrust (e.g., Visona et al., 2012; Searle, 2013, and references therein), or in

NU

transpressional settings (e.g., Gleizes et al., 1998; Iacopini et al., 2008; Archanjo et al., 2013). In the case of post-collisional extension, granitization is usually framed in a

MA

process of gravitational collapse (Coney and Harms, 1984; Dewey, 1988). Crustal melting in this setting is favored by heat input from cooling of mantle-derived magmas

D

and/or by uplift of the asthenosphere to compensate for lithosphere attenuation (England

PT E

and Thompson, 1984, 1986; Huppert and Sparks, 1988; Patiño Douce et al., 1990). Extensional faults atop the migmatised domains (e.g., Pico Sacro and

CE

Redondela-Beariz detachments; Díez Fernández et al., 2012b) do not have opposite sense of movement relative to Variscan thrusts in NW Iberia, which are characterized

AC

by E-directed tectonic transport (Marquínez García, 1984; Ribeiro et al., 1990; Martínez Catalán et al., 1996, 2002; Marcos and Farias, 1999; Díez Montes, 2007). Instead, those extensional detachments define a paired system of divergent detachments that moved to the NNW and SSE, i.e. following almost orthogonal vectors relative to previous thrusting (Díez Fernández et al., 2012b). Finally, the absence at the base of the migmatised domains of the Padrón dome of a major thrust, coeval with D2 deformation, does not support an extrusion model in NW Iberia for D2.

ACCEPTED MANUSCRIPT A transpressional setting, dominated by compression and featured by strike-slip shear zones, is not likely for the case of D2 in the Padrón dome either. All the strike-slip shear zones in NW Iberia are superimposed to a low-dipping main foliation (e.g., Iglesias Ponce de Leon and Choukroune, 1980; Jiménez Ontiveros and Hernández Enrile, 1983), which in the case of the Padrón dome it would correspond to S2. D2

PT

deformation is connected to extension-related structures (see discussion above), whereas

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strike-slip deformation is associated with D3 (Díez Fernández and Martínez Catalán,

SC

2012). Post-D2 exhumation of migmatised domains in a transpression-dominated setting have been suggested for other sections of Variscan Iberia (e.g., Central Iberia; Díez

NU

Fernández and Pereira, 2016), and it is probably the case in the study area. However, most of the exhumation path in Central Iberia was achieved via extension concentrated

MA

in migmatitic domes (Escuder Viruete et al., 1994, 1997). Similarly to the study area, extension-related structures in Central Iberia represent the second phase of deformation

D

at regional scale, and the age of superimposed (D3) strike-slip shear zones is equivalent

PT E

in both regions (ca. 310-305 Ma; Rodríguez et al., 2003; Gutiérrez-Alonso et al., 2015; Díez Fernández and Pereira, submitted).

CE

It is considered that the climax of crustal thickening during D1 led to a gravitationally and thermally imbalanced lithosphere. Such lithosphere experienced a

AC

process of thermal maturation and relaxation. Then it collapsed and flowed laterally to reduce gradients of gravitational potential energy (Alcock et al., 2009, 2015; Martínez Catalán et al., 2014). So far, there is no evidence for a large mantle contribution in the formation of the granitic rocks of the Padrón dome or in other similar Variscan granitic areas nearby (Gil Ibarguchi et al., 1984; Bellido et al., 1992). Therefore, the gravitational reequilibration at this point was probably achieved via progressive

ACCEPTED MANUSCRIPT lithosphere thinning, the uplifting of the asthenosphere being one of the main potential sources of external heat (if any).

6.2. Intermediate-D2 stage: recumbent folds A second stage in the exhumation of the relative autochthon was characterized

PT

by the development of recumbent folds of regional extent (Fig. 6c). These folds were

RI

nucleated at the expense of newly-formed mechanical instabilities, such as the

SC

increasing presence of irregular, sheet-like bodies of granitoid. Both the nucleation and amplification of these folds took place within a portion of the middle-upper crust

NU

subjected to decompression (e.g., growth of S2-andalusite after sillimanite). Top-to-theNNW kinematic criteria observed along the S2 (Fig. 3d; see section 4.2) and the

MA

stretching lineation associated with D2 suggest that tectonic flow at this stage was also highly oblique to the trend of the orogen.

D

Folding in migmatised crust is favored by numerous facts, including the

PT E

completely ductile behavior of the crust under partial-melting conditions, the existence of viscosity contrasts between melted and non-melted rock, and the availability of

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mechanical instabilities defined by the irregular shape of the multiple patches and larger intrusions of magma to nucleate folds from.

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Large-scale recumbent folds can be formed both in compressional or extensional settings. Tectonic models dominated by compression can be discarded for explaining the Portomouro synform, because D2 represents a stage related to extension. Moreover, neither the staircase thrust model conceived for the Helvetic nappes (Ramsay, 1981; Ramsay et al., 1983), nor the standard model of nappe folds (Hatcher and Hooper, 1992) fit the geology of the study area. Neither a discrete fault at the base of the

ACCEPTED MANUSCRIPT recumbent fold structure nor significant deformation gradients along the limbs of D2 folds have been identified. The formation of recumbent isoclinal folds within an extending lithosphere has been described for a number of cases (Platt, 1982; Froitzheim, 1992; Orozco et al., 1997; Harris et al., 2002; Arango et al., 2013). Following the model of Dietrich and

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Casey (1989), which was initially conceived for a compressional setting, a combination

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of pure shear and heterogeneous simple shear may explain the formation of fold nappes.

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The central role of shearing, and particularly non-coaxial strain (asymmetric S2 microstructures and development of some discrete shear bands), in the formation of this

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type of folds makes that model suitable for extensional settings. Accordingly, the relative autochthon could be equivalent to a broad, low-dipping ductile shear zone. This

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shear zone would have displaced horizontally an upper and colder crustal section (the allochthonous complexes) over a lower and hotter mid- to high-P granulitic crust

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(currently not exposed) (Díez Fernández et al., 2012b). Given the fold width of the

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Portomouro synform (Fig. 3b) and the penetrative nature of S2 in the whole study area, this virtual shear zone would have a minimum of 5km in thickness and represented a

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section of the crust subjected to widespread ductile deformation. Large-scale recumbent folding in these cases is favored by either incremental

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tilting of former planar anisotropies within the shear zone (e.g., early-S2; Arango et al., 2013), or by changes in the geometry of the shear zone boundaries (equivalent to the staircase thrust model of Ramsay et al., 1983). Both possibilities are very likely during the development of a migmatitic dome, where progressive or pulsed upwarping of its suprastructure is expected in response to the exhumation of its molten infrastructure (buoyant rise of the lower crust; Vanderhaeghe, 2009). A third possibility is that the boundaries of the granitic massif were oblique and dip steeper than the D2 shear planes

ACCEPTED MANUSCRIPT before recumbent folding (Froitzheim, 1992). Field observations suggest a markedly concordant character for the syn-kinematic granitoids, whose boundaries are usually sub-parallel to the low-dipping foliation. However, the rather irregular geometry of the massif of syn-kinematic granitoids after D2 (Fig. 3b) suggest a somewhat discordant nature of this massif before recumbent folding.

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The type of structures produced in migmatite domes changes depending on the

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rate of vertical crustal flow relative to subhorizontal flow (Whitney et al., 2004b).

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Dome-like structures (e.g., diapirs) rise when vertical flow dominates, whereas channels (either flat or inclined) emerge when lateral flow prevails. Migmatite nappes, on the

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other hand, result from a combination of vertical and lateral flow. These nappes may include a series of complex structures, such as cascades of folds, if vertical flow is

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slightly stronger, or recumbent folds (e.g., climbing folds) if the rate of subhorizontal flow is higher. The geometry and asymmetric microstructures of the foliation associated

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with the development of the Portomouro synform, and the formation of this fold itself,

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suggest dominant subhorizontal flow in the Padrón dome during the intermediate-D2 stage. This adds to the subhorizontal flow that existed during the early-D2 stage, needed

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to create the elongate shape of the granitic massifs before they were folded. All of this pictures a partially re-equilibrated crust still ruled by subhorizontal mass movements,

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even after significant melt production. Most of the granite-diatexite massif studied here does not preserve evidence for a pre-recumbent folding planar fabric. This suggests that cooling and crystallization of the granitic melts were not completed by the onset of D2 recumbent folding. Both of them must have taken place in the course of fold nucleation through amplification. A significant volume of granitic rocks shows fabrics axial planar to the Portomouro synform. These fabrics are either solid-state fabrics or magmatic fabrics (see section 3),

ACCEPTED MANUSCRIPT what suggests progressive cooling and crystallization during fold amplification. But there are also some zones of the granite-diatexite massif lacking of a fabric, what suggests that magma crystallization was not even complete after folding and/or that melt production/migration was still ongoing. The latter possibility connects with the intrusion of the late-kinematic two-mica granites, which lack of S2 but are affected by

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subsequent deformation (D3). Overall, the density contrast between the melts that were

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being folded and their host rock implies that there were some gravity forces held in the

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granites and able to generate a component of vertical flow within the Padrón dome at this stage (e.g., Vanderhaeghe, 2009), but apparently not strong enough to overcome the

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existing subhorizontal flow.

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6.3. Late-D2 stage: onset of late extensional detachments Deformation concentrated along discrete detachments during the final stage, thus

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producing further tectonic removal of the overlying tectonic pile (Fig. 6d). This stage is

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represented by the Pico Sacro detachment. Until now, the mechanics of both this detachment and the rocks at its footwall had been considered as reflecting the same

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extensional event (Gómez Barreiro et al., 2010; Díez Fernández et al., 2012b). The data presented here allow further constrains, and demonstrate that the Pico Sacro detachment

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is in fact a relatively late extensional event within D2, and follows a disparate structural evolution of the migmatitic dome that lies underneath. Accordingly, non-coaxial deformation concentrated in the upper crust at this stage, whereas widespread ductile deformation waned (declined?) after recumbent folding. Melt migration, on the other hand, did not cease, as bodies of granitic magma were still intruding the upper crust after the development of the Pico Sacro detachment.

ACCEPTED MANUSCRIPT D2 extensional deformation was replaced by a new phase dominated by subhorizontal contraction (D3; Fig. 6e). Upright folds and strike-slip shear zones (not represented here) modified the previous structure and contributed to a further exhumation of the relative autochthon. This type of structures are found almost all across the Iberian Massif (e.g., Iglesias Ponce de Leon and Choukroune, 1980; Ribeiro

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et al., 1990; Pereira et al., 2010; Pérez-Cáceres et al., 2015), so their origin is likely

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related to far-field stresses acting at the scale of the orogen, i.e. tectonic forces sourced

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from the collision of Gondwana and Laurussia by the end of the Variscan orogeny (e.g., Martínez Catalán, 2011). The role of equivalent tectonic forces in the course of D2 was

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discussed by Díez Fernández et al. (2012b), who argued that the orogen-oblique character of the subhorizontal extensional flow that characterizes some of the

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migmatite-cored domes of the Iberian Massif was favored by a strike-slip movement of Gondwana relative to Laurussia following the climax of their collision. Net orogenic

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flow in this context (D2) would result from a combination of tectonic and gravitational

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forces. The first ones would condition lateral mass flow to orogen-oblique vectors, whereas the second ones would be responsible for initiating and sustaining

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subhorizontal extensional flow.

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7. Conclusions

A syn-orogenic granitic massif located in the eastern flank of the Padrón migmatitic dome is deformed by a large-scale isoclinal recumbent fold, the Portomouro synform. The granitic rocks were hosted by partially-molten crust and represented subconcordant plutons (laccoliths) inferred to be stretched in a NNW-SSE direction before folding. Recumbent folding took place under general ductile shear directed to the NNW, and followed a main stage of crustal thickening (D1) and a subsequent stage of

ACCEPTED MANUSCRIPT thermal relaxation and gravitational collapse (early-D2). Subhorizontal orogenic flow dominated during both the onset of extensional collapse (early-D2) and then during its intermediate stages. Sheet-like granitic bodies intruded during early-D2, and km-scale recumbent folding took place during intermediate stages of D2. During late-D2, deformation concentrated along discrete extensional detachments that cut across

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previous structures.

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Our analysis suggests that the nucleation of recumbent folds in extending

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orogenic crust is plausible for the case of large-scale syn-orogenic granitoids as long as widespread ductile subhorizontal flow dominates after significant melt production. The

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whole set of mechanical anisotropies formed during previous stages of orogenic evolution, such as intrusions of syn-orogenic granitoids, would create a favorable

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rheological setting for the nucleation of folds, the geometry of which would depend on the geometry of existing crustal flow. In tectonic regimes dominated by extensional

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subhorizontal flow, recumbent folds seem the most likely structures to be formed.

8. Acknowledgments

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We thank Rodolfo Carosi and an anonymous reviewer for insightful comments and suggestions. Financial support has been provided by Instituto Geológico y Minero

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de España (Project IGME 2281). Rubén Díez Fernández appreciates financial support from Ministerio de Economía y Competitividad (Spain) through its Juan de la Cierva postdoctoral program (JCI-2012-11967). This work is a contribution to IGCP project 648 (Supercontinent Cycle and Global Geodynamics).

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Figure Captions

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Fig. 1. Simplified map of the Variscan orogen, after Martínez Catalán et al. (2007) and

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Díez Fernández and Arenas (2015).

Fig. 2. Simplified geological map and cross-section of the NW Iberian Massif (after

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Martínez Catalán et al., 2007; Díez Fernández et al., 2011). Location of the study area is indicated.

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Fig. 3. (a) Geological map and (b) cross-section of the eastern limb of the Padrón dome,

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focused on the Milladoiro antiform. (c) Stereoplot containing calculation of the Beta axis of D3 folds (F3) using S2 planes, Mean Principal Direction of D3 fold axes (F3) and Mean Foliation Plane for S3. (d) Stereoplot showing the trend of D2 fold axes (F2) and D2 stretching lineation (L2s). Reference Points or sites for understanding the deformation history of the study area are indicated as RP-X in the map and crosssection. UTM coordinates (Zone 29). Location of pictures shown in Figures 4 and 5 is included.

ACCEPTED MANUSCRIPT Fig. 4. (a) Main foliation (S2) in a non-migmatitic schist, with quartz, biotite, muscovite, sillimanite, and garnet. (b) Main foliation (S2) in the melanosome of a migmatitic metasedimentary rock. Note the growth of subhedral andalusite crosscutting fibrolitic sillimanite, and the disorganized growth of late muscovite. (c) Andalusite-bearing paragneiss with fibrolitic sillimanite enclosed within anhedral andalusite (whitish

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matrix). (d) D2 crenulation cleavage (S2) defined by quartz, biotite, muscovite, and the

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growth of subhedral and euhedral andalusite affecting a proto-S2 fabric containing

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quartz, biotite and sillimanite. (e) D3 subvertical crenulation cleavage (S3) affecting a migmatitic banding in metasedimentary rocks (S2). (f) Solid-state deformation features

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observed in the leucosomes of the migmatitic paragneisses (internal deformation and recrystallization, grain-size reduction, foliation anastomosing, myrmekite, and

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development of discrete shear zones). Note the presence of garnet in the leucosome.

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Geographical location is indicated in Figure 3b.

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Fig. 5. (a) Syn-kinematic granites showing penetrative D2 ductile deformation (S2). (b) Preferred mineral orientation in syn-kinematic granites. Mineral grains are affected by

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solid-state deformation (internal deformation and recrystallization, grain-size reduction, foliation anastomosing, and boudinage of competent minerals). (c) Subconcordant

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leucogranite lenses included in foliated migmatititc paragneisses (S2). (d) Tectonic banding (quartz-rich and mica-rich layers) crosscut by the main foliation (S2) in nonmigmatitic paragneisses (quartz, biotite, muscovite). This banding represents an earlyD2 foliation that suggests the existence of yet another previous mineral fabric (S1). (e) D2 recumbent folds affecting the migmatitic banding in the metasedimentary rocks. S2 is axial planar to these folds. (f) Mineral and stretching lineation (L2s) in foliated

ACCEPTED MANUSCRIPT paragneisses (S2). Lineation is defined by mica, and stretched quartz and feldspar. Geographical location is indicated in Figure 3b.

Fig. 6. Simplified tectonic evolution of the eastern flank of the Padrón migmatitic dome. (a) Initial stages of collision and crustal thickening (D1) led by the emplacement

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of the Allochthon onto its relative autochthon (Parautochthon and Autochthon). The

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sketch to the left is an oversimplification of the resulting tectonic stack. Note that the

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Basal and Ophiolitic allochthonous units are gathered into a single unit referred to as Allochthon. (b) Early stage of orogenic collapse (early-D2) dominated by partial

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melting, generation of large granitic massifs, and development of a shallow dipping foliation (proto-S2). (c) Intermediate stage of orogenic collapse (intermediate-D2)

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characterized by N-NW-directed extensional flow and the formation of recumbent folds and associated axial plane foliation (S2). (d) Late stage of orogenic collapse (Late-D2)

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featured by extensional detachments (top-to-the N-NW kinematics) cutting across

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previous D2 record. Details about the structural position of allochthonous units are included. (e) Onset of renewed crustal thickening (D3) and development of upright folds

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and subvertical crenulation cleavage (S3).

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Graphical abstract

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Exhumation of mainland Gondwana in the Iberian Massif resulted in lithosphere melting.

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Syn-orogenic granitic massifs feature the Variscan middle crust of NW Iberia. Lateral spreading of extending lithosphere favored nucleation of recumbent folds.

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Recumbent folds can be formed from granitoids in extension-dominated settings.

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Protracted subhorizontal flow distorts early extensional record in migmatitic domes.