Subduction- and exhumation-related structures preserved in metaserpentinites and associated metasediments from the Nevado–Filábride Complex (Betic Cordillera, SE Spain)

Tectonophysics 644–645 (2015) 40–57

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Subduction- and exhumation-related structures preserved in metaserpentinites and associated metasediments from the Nevado–Filábride Complex (Betic Cordillera, SE Spain) Antonio Jabaloy-Sánchez a,⁎, María Teresa Gómez-Pugnaire b,c, José Alberto Padrón-Navarta d, Vicente López Sánchez-Vizcaíno e, Carlos J. Garrido c a

Departamento de Geodinámica, Facultad de Ciencias, Universidad de Granada, Campus Fuentenueva, s/n 18002 Granada, Spain Departamento de Mineralogía y Petrología, Facultad de Ciencias, Universidad de Granada, Campus Fuentenueva, s/n 18002 Granada, Spain c Instituto Andaluz de Ciencias de la Tierra, CSIC-Universidad de Granada, Avda. de Las Palmeras no. 4, 18100 Armilla, Granada, Spain d Géosciences Montpellier, Univ. Montpellier 2 & CNRS, 34095 Montpellier, France e Departamento de Geología, Escuela Politécnica Superior, Universidad de Jaén (Unidad Asociada al CSIC-IACT Granada), Alfonso X el Sabio 28, 23700 Linares, Spain b

a r t i c l e

i n f o

Article history: Received 24 March 2014 Received in revised form 12 December 2014 Accepted 13 December 2014 Available online 8 January 2015 Keywords: Atg-serpentinite Chl-harzburgite Transpressional deformation Brittle deformations Betic Chain

a b s t r a c t The Cerro del Almirez massif (Nevado–Filábride Complex, Betic Cordillera, SE Spain) is composed of antigorite (Atg-) serpentinite and chlorite (Chl-) harzburgite separated by a thin reaction front formed in a palaeosubduction setting. These ultramafic rocks preserve unique prograde structures related to the pre- and peak high-pressure event (1.6–1.9 GPa and 680–710 °C). The oldest subduction-related structures are preserved in Atg-serpentinites: a penetrative S1 foliation and associated L1 stretching lineation that formed in a non-coaxial regime with a top-to-the-W sense of shear. This planar–linear fabric is crosscut by olivine ± Ti-clinohumite veins, formed during the prograde breakdown of brucite and pre-metamorphic clinopyroxene, which form a system of veins decimetres long. They record embrittlement processes due to the release of 6 vol.% of H2O associated with brucite dehydration. The growth of prograde olivine and/or tremolite porphyroblasts is syn- to postkinematic in relation to the S1 foliation. Further reactions at higher temperature related to the complete breakdown of the Atg (i.e. Atg-out) are post-kinematic to S1. Prograde Chl-harzburgite is crosscut by sets of conjugate zones associated with grain-size reduction of olivine grains. These grain-size reduction zones are interpreted as brittle structures generated by hydrofracturing due to overpressure fluids in a compressional setting at lowdifferential stresses. Structures related to the exhumation process are mainly preserved in the metasedimentary host rock, where an S2/L2 planar–linear fabric developed within a shear zone dominated by a non-coaxial regime with a top-to-the-W sense of movement in a transpressional regime. Peak metamorphic conditions deduced for the schists are similar in pressure (1.3–1.9 GPa) but lower in temperature (560–590 °C) compared to the ultramafic rocks in contact with them, suggesting a major shear zone at the base of the ultramafic massif during D2 deformation in a right-handed transpressional regime. © 2015 Elsevier B.V. All rights reserved.

1. Introduction Unravelling the structural and tectonic evolution of subducting slabs during the deformation related to prograde metamorphism is of capital importance to understanding the evolution of the rocks within a subduction zone in its early stages (i.e. at relatively shallow depths, around 40–80 km). However, structures related to these early stages of subduction and metamorphism are usually overprinted and obliterated during exhumation (e.g., Hermann et al., 2000) and, as a result, our knowledge of the deformation mechanisms of the different parts of the subducting ⁎ Corresponding author. Tel.: +34 958 243365; fax: 34 958 248527. E-mail address: [email protected] (A. Jabaloy-Sánchez).

http://dx.doi.org/10.1016/j.tecto.2014.12.022 0040-1951/© 2015 Elsevier B.V. All rights reserved.

slab mainly relies on numerical models (e.g., Angiboust et al., 2012; Chemenda et al., 1996; Cloetingh et al., 1999; Faccenda et al., 2008; Guillot et al., 2000, 2009) with only limited field constraints (Agard et al., 2009; Angiboust et al., 2012; Hermann et al., 2000; Padrón-Navarta et al., 2012). One approach to circumvent this problem is to investigate the deformation features of several lithologies with contrasting chemical compositions since the pressure and temperature (PT) during deformation can be better controlled due to the higher number of prograde reactions that can be used to track those conditions. Moreover, the structural analysis of a broader spectrum of lithologies also ensures a variety of rheological behaviour. Both features can be exploited to examine different segments of the pressure–temperature–time evolution in subduction settings.

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Ultramafic rocks are especially suitable to unravel subduction and exhumation structures because, despite the limited number of metamorphic reactions they undergo over a wide pressure and temperature range, these metamorphic reactions are very well constrained both from experimental (Ulmer and Trommsdorff, 1995) and from field studies (Hermann et al., 2000; Rebay et al., 2012; Scambelluri et al., 1995; Trommsdorff and Evans, 1974), and they usually occur in the subducting slabs associated with other lithologies such as metapelites and/or metabasites. The Nevado–Filábride Complex of the Betic Cordillera (Fig. 1) is of particular relevance to this task as it contains a continental, mainly metapelitic, sequence metamorphosed during the Alpine orogeny with associated serpentinite lenses metamorphosed at high-pressure (1.6–1.9 GPa) under the highest temperatures reported worldwide for an antigorite-bearing (Atg) serpentinite (680 °C) in this pressure range (Padrón-Navarta et al., 2010b). Atg-serpentinite underwent further heating (up to 710 °C) that resulted in its dehydration and the formation of its prograde counterpart, chlorite- (Chl-) harzburgites (Padrón-Navarta et al., 2011; Puga et al., 1999; Trommsdorff et al., 1998). These conditions were attained during the Middle Miocene subduction of several continental units from the Nevado–Filábride Complex (Gómez-Pugnaire et al., 2004, 2012; López Sánchez-Vizcaíno et al., 2001; Platt et al., 2006). The main aims of this work are as follows: (1) to describe structures that developed during progressive burying in the subduction zone of a metaserpentinite lens going beyond the antigorite dehydration reaction, (2) to describe structures explaining the subsequent emplacement of the ultramafic rocks on the top of a continental lithological sequence, the Nevado–Filábride Complex, and (3) to establish the deformation regime of a subducting slab both in the early stages of subduction and during the decoupling processes taking place in that slab at relatively low depths above 70 km.

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2. Geological setting The western Mediterranean region includes several different segments of the Alpine orogen: the Betic–Rif belt, the Tell, the Apennines, the Alps, and the Pyrenees–Basque–Cantabrian Mountains (e.g. Jolivet and Faccenna, 2000; Royden, 1993). Orogenic processes began during the Cretaceous and involved several episodes of high-pressure/lowtemperature (HP/LT) metamorphism that may correspond to collisions between continental ribbons and microplates that existed in the area between the converging plates of Africa and Europe (e.g., Dercourt et al., 1986; Stampfli, 2000). The Betic Cordillera (southern Spain) (Fig. 1) is part of the Betic–Rif belt and formed as a result of the collision of an allochthonous terrain (Alborán Domain, Balanyá and García-Dueñas, 1987) with the southern palaeomargin of the Iberian plate during the Neogene (Balanyá and García-Dueñas, 1987; Platt et al., 2013). The Alborán Domain (Fig. 1a) includes several tectonic complexes, including the uppermost Maláguide Complex. It is composed of a Paleozoic basement (deformed and metamorphosed at green-schist facies conditions during the Variscan orogeny) that is overlain by a very weakly metamorphosed Mesozoic–Paleogene sedimentary cover (e.g. Fernández-Fernández et al., 2007). The Alpujárride Complex crops out structurally below the Maláguide Complex and consists of Paleozoic to Triassic rocks metamorphosed at high pressure/low temperature (e.g., Goffé et al., 1989; Tubia and Gil-lbarguchi, 1991) at 23–21 Ma (Janots et al., 2006; Sánchez-Rodríguez and Gebauer, 2000). Alpujárride rocks clasts appear within the unconformable Lower–Middle Miocene sedimentary formations covering the Alborán Domain (e.g. Fernández-Fernández et al., 2007), implying that this complex was completely exhumed at that time.

a

b

c Fig. 1. a) Location of the study area within the Betic–Rif orogenic system. b) Geological map of the southwestern Betic Chain. Location of the study area in Fig. 2 is marked with a blue rectangle. c) Lithological sequence of the Nevado–Filábride Complex in the Cerro del Almirez area. Ages according to the references indicated by the numbers in brackets. b, modified from Vázquez et al. (2011).

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The deepest Nevado–Filábride Complex (affected by high-pressure/ intermediate-temperature metamorphism and cropping out in the core of three east–west trending antiforms) is separated from the overlying Alpujárride Complex by a ductile (1–2 km thick) to brittle low-angle normal fault that formed during late exhumation processes (Agard et al., 2011; Galindo-Zaldívar et al., 1989; González-Lodeiro et al., 1984; Jabaloy et al., 1993) (Fig. 1a). The ultramafic rocks that are the subject of this paper are located at the top of the Nevado–Filábride Complex (Fig. 1b). Most current tectonic interpretations of the Betic Cordillera are the result of the recent determination of a Miocene age for the Alpine metamorphism of the Nevado–Filábride Complex (Behr and Platt, 2012; Gómez-Pugnaire et al., 2004, 2012; López Sánchez-Vizcaíno et al., 2001; Platt et al., 2006). This age is much younger than in previous, mainly Ar/Ar results (i.e., Augier et al., 2005c; De Jong et al., 1992; Monié et al., 1991), as well as the age of metamorphism of the overlying Alpujárride Complex (Janots et al., 2006; Sánchez-Rodríguez and Gebauer, 2000). These data indicate that the Alpujárride Complex was already exhumed at the time the Nevado–Filábride rocks underwent high-pressure metamorphism (López Sánchez-Vizcaíno et al., 2001; Platt et al., 2006). This has led to the redefinition of the Alborán Domain as only the two uppermost complexes (Alpujárride and Maláguide). The Nevado–Filábride Complex is now considered part of the Iberian Massif palaeomargin that was subducted below the allochthonous Alborán Domain just after the subduction of the Alpujárride Complex at 20 Ma, and accreted below the Alpujárride rocks during a late exhumation event (Augier et al., 2013; Behr and Platt, 2012; Booth-Rea et al., 2007, 2012; Gómez-Pugnaire et al., 2004, 2012; Platt et al., 2013). In this context, the structural evolution documented in this paper was related to the initiation of a subduction zone rather than the initial fabrics that develop within a subducted slab entering a subduction zone that may have been active for some period of time. 2.1. Lithological and tectonic features of the Nevado–Filábride Complex 2.1.1. Lithology The metasedimentary sequence of the Nevado–Filábride Complex consists, from bottom to top (Fig. 1c), of the following rock formations according to Gómez-Pugnaire et al. (2012): 1) A thick and monotonous sequence (more than 2000 m thick) of lowgrade micaschists (Veleta Unit, Puga et al., 2002) dated as Neoproterozoic (Gómez-Pugnaire et al., 1982) to Devonian (Laborda-López et al., 2013; Lafuste and Pavillon, 1976). 2) A similar lithological sequence of graphite-bearing micaschists (Montenegro Fm; Fig. 1b, c), although with a greater degree of metamorphic recrystallization (Augier et al., 2005a,c; Gómez Pugnaire and Franz, 1988; Puga et al., 1975) than those of the Veleta Unit rocks. 3) A sequence of fine-grained, feldspar-bearing micaschists with numerous quartzite intercalations (up to 1000 m thick, Tahal Fm in Fig. 1b, c). Scarce staurolite–kyanite-bearing micaschists and marble intercalations (up to 1 m thick) and orthogneiss and metabasite bodies may also appear in the upper part of this formation (e.g., Vissers, 1981). 4) A marble and calcschist sequence (Huertecica Marbles, see Nijhuis, 1964), mostly consisting of alternating calcschists with marble layers of extremely variable thickness, that pass gradually into very thick massive marbles with intercalations of garnet– chloritoid–(kyanite-) micaschists. At the base of this formation are discontinuous and thin metaevaporite layers (Gómez-Pugnaire et al., 1994). Interbedded metabasite and serpentinite layers are common. 5) A heterogeneous upper sequence (Casas and Nevada Fms, according to Kampschuur, 1975) consisting of a thick layer of white massive marbles (up to 20 m in thickness) overlain by a sequence of

alternating orthogneisses, dark (graphite-bearing) and light garnet–kyanite–chloritoid micaschists, locally with staurolite, grey and yellow banded marbles (up to 2 m thick), and scarce metabasite and serpentinite bodies. The Nevado–Filábride lithological sequence is Paleozoic or older in age (Gómez-Pugnaire et al., 2004, 2012) and, based on U–Pb SHRIMP dating on zircons sampled in orthogneiss from the three upper formations, it could be interpreted as a Variscan basement that can be correlated with the Central Iberian Zone of the Iberian Massif. In the study area (Cerro del Almirez peak area, Sierra Nevada; Fig. 1b), the two lower Veleta and Montenegro formations (corresponding to the Veleta and Mulhacén tectonic units, respectively; Fig. 1c) are well exposed, but the uppermost Tahal Fm has been significantly reduced in thickness (maximum 200 m; Figs. 1a–4, e.g. MartínezMartínez et al., 2002) and only the upper light-coloured micaschists and thin marble layers are locally found (blue and yellow, respectively, in Figs. 2 and 4). The ultramafic rocks, which are the main focus of this study, appear in this area on the top of the metasedimentary sequence (Figs. 1c–4). They crop out in a ~400 m thick and ~2.3 km2 wide tabular body in the Cerro del Almirez peak and in several smaller, discontinuous ultramafic bodies scattered west of the main outcrop (Jansen, 1936) (Fig. 2). They are bounded by a lower contact with metapelites (Montenegro and Tahal Fms) and with discontinuous, often brecciated, tabular bodies (usually 10 m thick) of pale yellow marbles with abundant tremolite and, more rarely, with thin-bedded, well-layered white marbles that we attribute to the Casas Formation (Figs. 1c–4). The Cerro del Almirez ultramafic massif is composed of two main lithologies: (1) a circa 200 m-thick sequence of well-foliated Atgserpentinite in the structurally upper half of the massif, and (2) a circa 200 m-thick layer of more massive Chl-harzburgite in the lower half (Trommsdorff et al., 1998) (Figs. 1c, 3, and 4). Atg-serpentinite, mainly consisting of antigorite, olivine, and magnetite, contains subordinate clinopyroxene–tremolite (Cpx–Tr)-rich serpentinite layers (Trommsdorff et al., 1998) and scarce olivine ± Ti-clinohumite veins (López Sánchez-Vizcaíno et al., 2009). Small lenses of Si-rich (orthopyroxene-bearing) serpentinite are enclosed within the Chlharzburgite (see below, Padrón-Navarta et al., 2010b). Both Atgserpentinite and Chl-harzburgite contain metamorphosed rodingitized (i.e. mafic dykes with intraplate to MORB affinities, Schönbächler, 1999), clinopyroxenite layers, and dunitic pods (López SánchezVizcaíno et al., 2001; Puga et al., 1999; Trommsdorff et al., 1998).

2.1.2. Tectonic features The main regional structure affecting the upper part of the Nevado– Filábride sequence is a planar–linear fabric (S2/L2) produced by heterogeneous simple shear and composed of a foliation with associated stretching lineation. A top-to-the-west sense of movement of the simple shear component of the strain is indicated by quartz C axis preferred orientation fabrics, S–C structures (according Berthé et al., 1979a,b; Passchier and Trouw, 2005), rotated porphyroblasts, and asymmetric pressure shadows (Agard et al., 2011; Augier et al., 2005b; Galindo-Zaldívar et al., 1989; Jabaloy et al., 1993; Platt and Behrmann, 1986). The S2/L2 planar–linear fabric develops inside a ductile major shear zone (1–2 km thick) culminating in the low-angle normal fault acting as the Alpujárride/Nevado–Filábride detachment. This major shear zone has been interpreted as having an extensional character and being related with the late exhumation and unroofing of the Nevado– Filábride Complex (e.g., Augier et al., 2005b; Galindo-Zaldívar et al., 1989; Jabaloy et al., 1993). Nevertheless, it has recently been reinterpreted as produced in a compressional setting when the rocks of the complex were being uplifted within the subduction channel (Behr and Platt, 2013).

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a

b

c

Fig. 2. a) Geological map of the Cerro del Almirez area. See Fig. 1 for location. Below, the diagrams represent the stereographic projections of the orientation of the main S2 foliation (b) and the L2 stretching lineation (c) in the metasediments of the study area. Eigenvectors and their eigenvalues have been determined using the Stero32 software developed by K. Röller and C. A. Trepman from the Institute of Geology, Mineralogy and Geophysics of the Rhur University (Bochum, Germany).

The planar–linear fabric (S2/L2) was deformed by (F3) folds varying from millimetric to hectometric in size that developed a crenulation cleavage (S3) affecting the whole Nevado–Filábride sequence. Late ductile–brittle and brittle extensional crenulation cleavage, more intense towards the top of the complex, overprinted all previous ductile structures with a top-to-the-west-southwest sense of movement (Augier et al., 2005b; Galindo-Zaldívar et al., 1989; Jabaloy et al., 1993; Platt and Behrmann, 1986). 2.1.3. Metamorphic evolution Peak metamorphic assemblages in metasediments and metamorphosed mafic and felsic igneous rocks of the upper Nevado–Filábride Unit indicate P–T conditions ranging between 1.4 and 2.0 GPa and 600 and 690 °C (Gómez-Pugnaire et al., 1994; Nieto, 1996; Puga and Díaz de Federico, 1976). HP conditions were followed by retrograde conditions at 600–300 °C and 0.6–0.3 GPa (Augier et al., 2005c; Gómez-Pugnaire et al., 1994; Puga and Díaz de Federico, 1976). Experiments by Padrón-Navarta et al. (2010b) further constrain the maximum stability of antigorite (680 °C at 1.9 GPa) and the peak metamorphic conditions of Chl-harzburgite (680–710 °C and 1.6–1.9 GPa).

HP metamorphism occurred between 18 and 15 Ma during the Middle Miocene, as revealed by SHRIMP U–Pb zircon and Lu–Hf garnet ages (Gómez-Pugnaire et al., 2004, 2012; López Sánchez-Vizcaíno et al., 2001; Platt et al., 2006). 40Ar/39Ar ages range from 48.4 ± 2.2 Ma to 24.6 ± 3.6 Ma in amphiboles from the western Sierra de los Filabres (Monié et al., 1991), and from 43.57 ± 4.2 Ma to 9.62 ± 0.2 Ma in micas from Sierra de los Filabres (e.g., Andriessen et al., 1991; Augier et al., 2005c; De Jong, 2003; Hebeda et al., 1980; Monié et al., 1991; Platt et al., 2005, 2006). However, Andriessen et al. (1991), De Jong (2003), Hebeda et al. (1980), and Platt et al. (2006) demonstrate that there is excess argon in the Nevado–Filábride rocks and the K/Ar and 40 Ar/39Ar ages are older than the ages obtained by other independent radiometric methods. 3. Main structures of the ultramafic rocks 3.1. Atg-serpentinite In the few outcrops where the original compositional layering of the ultramafic rocks can be observed, it is defined by pyroxenite layers with

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5080

5076

Atg-serpentinite/Debris

a

Chl-harzburgite/Debris

n=2

Marble

41050

n = 11

Schist (undifferentiated)

n = 10 n=5

Atg-out isograd (~680 ºC)

n = 11

Mechanic contact n=7

NNE-SSW late synform in Atg-serpentinite

n = 20 n = 15

WNW-ESE minor F1 folds in Atg-serpentinite

n=5

41046

n = 14

b

Layering N

n=6 n=9 n = 14 n = 12

41042

n = 32

n=7

Poles of the S1 foliation

c

n=6

N

41038

n = 10

0

L1 lineation

d

N

100

n = 189 Eigenvector

200m

Poles of the GSRZ surfaces

e

N

GSRZ surfaces N

f σ3

σ2 σ1 n = 29 Eigenvector

n = 107 Eigenvector

Eigenvector Maximum density = 23.0

Fig. 3. a) Geological map of the Cerro del Almirez Massif (see Fig. 2 for location) and stereographic projections of the orientation of the main structures observed within the ultramafic rocks. b) Orientation of the layering defined by pyroxenite layers in Atg-serpentinite. c and d) Orientation of the main S1 foliation and L1 stretching lineation, respectively, in Atgserpentinite. e and f) Orientation of the GSRZ in Chl-harzburgite with the location of the stress axes (f). a, modified from Padrón-Navarta et al. (2010a).

a disperse orientation due to late folding ranging between N20°E and N80°E strikes and 25°SE to 90° dips (Figs. 3b and 5a). Foliated Atg-serpentinite is characterized by planar (S1) and planar–linear fabrics (S1/L1) (Table 1, Figs. 3 and 5a, b). In both fabrics, S1 foliation, marked by the orientation of antigorite (Fig. 5c and d), is strongly penetrative and usually displays low dips with an average subhorizontal orientation (Figs. 3a–d, 4 and 5a, b). On S1 surfaces (Fig. 5e), magnetite + chlorite aggregates have elongated shapes that define a stretching lineation (L1) (Figs. 3 and 5c). The high variability in L1 orientations is studied by means of the eigenvectors (v1, v2, v3) and their eigenvalues (E1, E2, E3, where E1 ≥ E2 ≥ E3) (e.g., Woodcock, 1977) (Fig. 3d). We have also determined the strength parameter C (C = ln (E1/E3)) proposed by Woodcock (1977), which ranges between 0 and infinity and measures the fabric

strength, with higher values of C indicating strong preferred orientations. L1 is distributed in a girdle following the mean low-dipping S1 foliation with two maxima: one, with two-thirds of the data, plots around a N64°E average strike and another, with around one-third of the data, plots around a N147°E strike. The strength parameter C has a low value (C = 1.637) in agreement with the presence of two strong preferred high-angle orientations. Small shear zones (C shear bands according to Berthé et al., 1979a,b; Passchier and Trouw, 2005) define S–C structures indicating a top-tothe-west sense of shear. Metarodingite dykes and diopside + tremolite aggregates included in Atg-serpentinite are also sheared and have a σ-shape with a top-to-the-WSW sense of shear (Fig. 5d and f). Several lines of evidence, from field to microscopic observations, support that the S1/L1 fabric formed during the prograde metamorphic

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a

45

1

1’ Almirez

NNE

SSW S1

2.400

Atg-out isograd

2.200

Landslide 2.000

b

S2

2

2’

NNE

SSW

2.400 2.200 2.000

c

3

3’

NW

S2

SE

2.400 2.200

S2

2.000

d

4

4’ SE

NW 2.100 2.000

S2

e

5’

5

Almirez

ENE

S1

2.400

WSW

Atg-out isograd

S2

Landslide

2.200 2.000 1.800

f

Fig. 4. a–e) Geological cross-sections of the Cerro del Almirez area. See Fig. 2 for location. Dashed green lines represent the axial traces of the F3 folds. f) Block diagram illustrating the geometry and relationships between the major structures described in the text. Dashed green lines represent the axial traces of the F3 folds and thick dashed black lines mark the prolongation of major tectonic contacts.

evolution undergone by Atg-serpentinite before reaching the metamorphic peak: 1) The layers of metamorphic diopside + Ti-Clinohumite (López Sánchez-Vizcaíno et al., 2005) show an internal planar–linear fabric coherent with the S1 in the Atg-serpentinite. 2) The occurrence of σ-type porphyroclasts of metamorphic diopside within the Atg-serpentinite (Fig. 5d, see also Padrón-Navarta et al., 2012, their Fig. 1b and c).

3) S1 foliation is overgrown by xenomorphic to prismatic idiomorphic porphyroblasts of olivine, most commonly below 1 cm but locally reaching up to 10 cm in length (Fig. 5g). Textural relationships at thin-section scale indicate that oriented growth of antigorite blades defining the S1 foliation also took place simultaneously or after crystallization of olivine grains (Padrón-Navarta et al., 2008, their Fig. 2a). 4) Tremolite appears in two different textural positions: first, as small grains in wedge-shaped strain shadows rimming pre-metamorphic

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A. Jabaloy-Sánchez et al. / Tectonophysics 644–645 (2015) 40–57

a

b

S1

layering

S1

500 μm

Atg

c

500 μm

d

S1 Tr

Cpx

S1 Atg

e

f W-E

L1 blackwalls

Ol

S1 C

g

500 μm

h Atg Tr S1

Tr Ol

S1 surface

Fig. 5. Occurrence of S1 schistosity in the Cerro del Almirez Atg-serpentinites. a) Relationship between a dark pyroxenite-layer and the later S1 schistosity. b) Typical outcrop of Atgserpentinite displaying penetrative S1 schistosity. c) Thin section picture of Atg-serpentinite with anastomosed S1 schistosity. S1 is also defined by almond-shaped aggregates of finegrained antigorite. d) Atg-serpentinite with deformed porphyroclasts of metamorphic clinopyroxene surrounded by S1. e) L1 stretching lineation defined by oriented magnetite and chlorite aggregates on the S1 surface. f) Section parallel to the stretching lineation in a sheared metarodingite dyke (pinkish) and black-wall (green) within Atg-serpentinite. g) View normal to the S1 surface cut by the growth of idiomorphic brown olivine (Ol) porphyroblasts and aggregates of whitish tremolite (Tr) crystals. h) Idiomorphic tremolite crystal (Tr) crosscutting the S1 foliation affecting antigorite. Abbreviations for mineral names after Whitney and Evans (2010).

clinopyroxene porphyroclasts (partially recrystallized to diopside). In the second case, idiomorphic tremolite porphyroblasts, up to 4 cm long and commonly associated with large olivine porphyroblasts clearly crosscut the antigorite S1 foliation (Fig. 5g and h). In these rocks, the prograde tremolite formation reaction (diopside + antigorite → olivine + tremolite + H2O) has been calculated to occur at temperatures above 600 °C in a pressure range of 1.7–1.9 GPa (López Sánchez-Vizcaíno et al., 2009). 5) S1 foliation can be crosscut by olivine ± Ti-clinohumite veins (Fig. 6a–d). Olivine crystals in the veins can be fibrous and have

elongated habits oriented normal to the walls (Fig. 6b). Angular irregularities in sharp vein walls that match across the vein (Fig. 6c), and curved vein tips (Fig. 6d) that may split into sharp-edged splays have also been observed. All these observations indicate brittle fracturing of the Atg-serpentinite during the formation of the olivine ± Ti-clinohumite mineral association. Several, asymmetrically folded veins (Fig. 6e) suggest that ductile deformation continued synchronously with the vein formation. 6) Olivine occurs as well as the main mineral filling metre-long joints with associated drag folds (hybrid joints in Hancock, 1985, and

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Table 1 Sketch of correlation between the different structures observed in the various lithologies of the Cerro del Almirez area and their relationships with the metamorphism. We include a proposal of deformational stages. The structures marked in bold letters and with * correspond to structures generated in a non-coaxial regime with a top-to-the-west sense of shear. Atg-serpentinites

Chl-harzburgites

Metasediments (Mulhacén Unit)

Metasediments Deformational Relationships with the (Veleta Unit) stage metamorphism

S1 (only observed in thin sections within D1 Prograde metamorphism S1/L1 (cut by brittle veins Residual S1/L1 (only visible near the and hybrid joints)⁎ Atg-out isograd and in scarce outcrops) microlithons and small outcrops) – GSRZ – – – Peak metamorphic conditions – D2 Peak to retrograde metamorphism – – S2/L2⁎ – D3 Retrograde metamorphism – – F3 folds D4 Retrograde metamorphism – – – Ssh/Lsh⁎ Late deformational stages common for all the lithologies, except the Chl-harzburgites (F4 folds, extensional crenulations cleavages, open folds and joints)

Mandl, 2005) that cut the main S1 foliation of Atg-serpentinite (Fig. 6f).

3.2. The contact between Atg-serpentinite and Chl-harzburgite The contact between Atg-serpentinite and Chl-harzburgite, very well exposed in the largest body of the Cerro del Almirez (Figs. 2–4), provides significant information for the timing of the S1/L1 fabric in the Atg-serpentinite. The Chl-harzburgite has been interpreted as being the result of Atgserpentinite dehydration due to the antigorite breakdown reaction (Trommsdorff et al., 1998). S1 and S1/L1 fabrics in Atg-serpentinite are overgrown by both the Atg-out isograd (Fig. 3a) and the breakdown products of the Atg-serpentinite (i.e., Chl-harzburgite), indicating that the S1 and S1/L1 fabrics are the oldest deformational structures in the whole ultramafic massif. Four lines of evidence support this interpretation: 1) The contact between Atg-serpentinite and Chl-harzburgite is very sharp and sub-horizontal at the scale of the massif (Padrón-Navarta et al., 2008, 2010a), whereas the foliation has steeper dips (20–30°) and defines two open WNW–ESE trending folds (one antiform and one synform) that are transected in both limbs by the antigorite-out dehydration isograd (Padrón-Navarta et al., 2010b; Fig. 3). This indicates that the two structures (S1 foliation and WNW–ESE folds) formed prior to the Chl-harzburgite (Fig. 3a). 2) In the contact between Atg-serpentinite and Chl-harzburgite, a narrow band of prograde lithologies occurs between Atg-serpentinite and the highest-degree Chl-harzburgite (Padrón-Navarta et al., 2011), and they clearly cut the S1 foliation at an angle of 20° to 30°. 3) Small lenses (up to ~ 3 m long and 2 m thick) of serpentinites are metastably preserved within the Chl-harzburgites (Padrón-Navarta et al., 2010b, 2011). They display an S1 penetrative foliation (N60°E strike and 33° dip towards the SW) sub-parallel to that of the overlying Atg-serpentinite and traversing the contact with Chlharzburgite. The S1 foliation in some of these serpentinite lenses is overgrown by sub-horizontal veins filled with olivine and orthopyroxene as well as by large olivine and orthopyroxene porphyroblasts formed during extensive static recrystallization that took place during the prograde path at temperatures above 580 °C for a pressure of 1.8 GPa (Padrón-Navarta et al., 2010a). These relationships indicate that the S1 is older than the olivine– orthopyroxene recrystallization stage. 4) A few metres away from the serpentinite lenses, Chl-harzburgites, regardless of their texture (granofelsic or spinifex-like as defined by Padrón-Navarta et al., 2010b; see next section), appear as generally undeformed rocks (see below), thus indicating complete overgrowth and obliteration of previous fabrics by the newly formed mineral assemblages. However, planar and elongated magnetite and/or chlorite aggregates may locally preserve a residual S1/L1 fabric (Fig. 6g).

3.3. Chl-harzburgite Prograde Chl-harzburgite includes rocks with granofels (PadrónNavarta et al., 2011) and spinifex-like (Trommsdorff et al., 1998) textures. Granofels Chl-harzburgite is characterized by coarse granular olivine, chlorite flakes, and prismatic orthopyroxene with an interlocked texture. It is intercalated with the spinifex-like rocks at the metre to tens-of-metres scale throughout the entire massif (Padrón-Navarta et al., 2011). The spinifex-like texture is characterized by the growth of large (up to 12 cm in length), arborescent olivine crystals (Fig. 6g, h) in a matrix composed of radial aggregates of acicular orthopyroxene crystals of variable size, long chlorite flakes (b1 mm), aggregates of magnetite, and rare idiomorphic tremolite crystals. There is no general orientation to the spinifex-like texture at the massif scale although preferred orientations exist locally (Padrón-Navarta et al., 2010a). Grain size and the preferred orientation of olivine and orthopyroxene crystals in spinifex-like Chl-harzburgite vary at the centimetre to metre scale. Both the spinifex-like and the granofelsic textures are partially overprinted by grain-size reduction domains a few millimetres to several metres thick (Fig. 6h). Grain-size reduction zones (GSRZ, Table 1) are the only deformational mesostructures affecting the Chl-harzburgite. Thick GSRZ clearly have more steeply-dipping orientations (30–60°) than the foliation in Atg-serpentinite (Fig. 3c, e, and f). Analysis of pole orientation data for 107 GSRZ (regardless of their thickness) shows a girdle distribution roughly following a N–S surface that dips around 70° towards the west (Fig. 3e and f). In the girdle, two major clusters of GSRZ poles plot at 40° to each other, indicating that most surfaces belong to one of two conjugate sets of planes: one with a mean N70°E strike and 40° dip towards the SSE, and the other with a mean N–S strike and 30° dip towards the E (Fig. 3e and f). A third cluster with fewer poles has a mean E–W strike and dips 40° towards the N.

3.4. Relationships with the metasediments Ultramafic bodies in the study area occur on top of the Nevado– Filábride metasedimentary formations, either the graphitic micaschists from the Montenegro Fm or the upper, light-coloured micaschists from the Tahal Fm (Figs. 2–4). In the northernmost contact of the Cerro del Almirez body (yellow in Fig. 3) and in the westernmost ultramafic body, thin, discontinuous, tabular bodies of marbles occur in the contact between metapelites and ultramafic rocks. In most cases, prograde Chl-harzburgites overlie either graphitic or light schist or yellow marbles. The contact surface is parallel to the S2 foliation of the lower rocks, but Chl-harzburgite itself (with either spinifex-like or granofelsic texture) lacks any internal deformation near the contact. Only in a very few localities (northern part of the Cerro del Almirez body) can Atg-serpentinites be found in contact with wall rocks and, specifically in the outcrops mentioned, in contact with light metapelites (Tahal Fm, Fig. 2). In every case, wall-rock metasediments display the same S2/L2 planar–linear fabric as the main structure (Table 1, Fig. 7a): clustered

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a

b

S1 Ti-Chu fibrous olivine

Ol

olivine porphyroblasts

c Matching vein walls

Curved vein tip

d Ol

Ol

e

f Drag folds

Ol Ol

g

h

S1 /L1 GSRZ GSRZ’ Opx Ol Spinifex-like texture

Mag+Chl

Spinifex-like texture

Fig. 6. Field occurrence of the main structural features observed in the ultramafic rocks of the Cerro del Almirez area. a) Olivine + Ti-clinohumite vein crosscutting the antigorite S1 foliation. b) Vein filled with elongated fibrous olivine crystals indicating movement normal to the wall. Width of the vein around 2 cm. c) Curved vein tip at the end of a sharp olivine vein. d) Angular irregularities in sharp olivine vein walls matching across the vein and indicating movement both parallel and normal to the walls (characteristic of hybrid joints). e) Olivine vein deformed by asymmetrical folds. f) Small, left-handed, faulted hybrid joint filled with olivine; fold drag can be observed in the two blocks. g) Residual S1/L1 planar–linear fabric marked by oriented magnetite and chlorite aggregates within Chl-harzburgite with spinifex-like texture. h) GSRZ and GSRZ′ (green) cutting the spinifex-like texture of Chl-harzburgite. Abbreviations for mineral names after Whitney and Evans (2010).

penetrative S2 schistosity with a mean N50°E strike and a dip of 20° towards the SE (Fig. 2b), and stretching lineation (L2) defined by elongated aggregates of the minerals described in Section 4 (Figs. 2b, c and 7c) and clustered around a mean N80°E trend and 14° dip towards the WSW (Fig. 2c). The S2/L2 planar–linear fabric has associated mm- to cm-thick shear bands (C-bands) at low angles to S2 that define standard S–C structures (Fig. 7d) with a top-to-the-west sense of movement. Quartz segregates in the schists are affected by isoclinal folds (F2) with fold axes usually

parallel to the L2 stretching lineation and affected by the C surfaces, thus defining σ structures with a top-to-the-west sense of shear. Sheath folds are scarce and most are isoclinal folds with straight hinges at low angles to the stretching lineation (Fig. 7c). In the study area, all the metasediments and the Chl-harzburgite bodies (Figs. 2 and 4) are affected by close to tight N-vergent F3 folds with wavelengths ranging between 800 m and 1 km (Table 1). F3 folds are cut by a shear zone at the base of the Mulhacén unit that, in turn, is slightly affected by N–S very open folds with immersion

A. Jabaloy-Sánchez et al. / Tectonophysics 644–645 (2015) 40–57

49

Q-domain S1

S2

Ph + Pg

500 μm

S2 C Ph

a

M-domain

C’

fold F2 hinge

L2

b

S2

C

C W-E

d

c Cld

Qz

80

2´

Qz

Ilm mantle

70

Ilm

20

2 Ph

g

Cld Ilm

Ph

mantle

e

Ap

10

Chl

0

Cld Alm Grs Sps Prp

15

Cld

10

Ph

0

h

f

5 0

10

20

Grt

500 μ 30

0.25

40

50

i

0.24 0.23 1’

0.22

Cld 1

500 μ

X Mg

0

4

8 12 16 20 24 28 32 36

Fig. 7. Textures and ductile mesostructures in the metasediments of the Cerro del Almirez area. a) Outcrop of dark schists of the Mulhacén Unit with the S1 crenulated foliation and the new S2 foliation. b) Thin-section of the quartz and phengite-rich domains (Q- and M-domains, respectively) defining the S2 foliation in the Tahal schists. Within the M-domains, large phengite porphyroblasts with decusate textures overgrow small phengite + paragonite mica grains that are only locally preserved. c) Quartz vein in the dark schists of the Mulhacén Unit with the L2 stretching lineation at low angles of F2 hinges. d) S–C and S–C′ structures (Berthé et al., 1979a,b; Passchier and Trouw, 2005), with a top-to-the-west sense of shear, developed during the D2 deformation phase in the dark schists of the Mulhacén Unit. e to i) Photographs and compositional profiles of garnet and chloritoid porphyroblasts of the Tahal schists (see Fig. 2 for location, Table 2 for mineral compositions). e) Garnet porphyroblast with indication of the analysed profile shown in f and below it a fractured chloritoid porphyroblast with an open joint filled by chlorite and white mica. g) Fragments of a garnet porphyroblast partially enclosed by chloritoid. Note the garnet fractures also filled by chloritoid. h) Snowball garnet porphyroblast with inclusion-free mantle partially enclosed by a chloritoid porphyroblast. The analysed profile shown in i is indicated in the chloritoid porphyroblast below. Abbreviations for mineral names after Whitney and Evans (2010).

towards the south. The Mecina extensional detachment (GalindoZaldívar et al., 1989) that separates the Nevado–Filábride Complex from the Alpujárride rocks is located around 300 m above the described outcrops.

4. Metamorphic conditions of wall-rock metasediments In the previous sections, it has been demonstrated that in the Cerro del Almirez area ultramafic rocks are located on top of a

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metasedimentary sequence of continental origin. To elucidate the origin of this structural anomaly, we compared the PT conditions and structures of the ultramafic rocks and underlying metasediments in order to assess whether they were in the current position prior to metamorphism or, conversely, their present relationships are the result of postmetamorphic exhumation. We estimated for the first time in this area the metamorphic conditions of the light-coloured schists from the Tahal Fm directly below the ultramafic rocks (Fig. 2), which occur in the upper part of the metasedimentary sequence and most probably record the metamorphic conditions corresponding to the formation of the above-described F1 folds, in whose synclines they occur (Fig. 4b–d). These rocks also record the metamorphic conditions attained during the later development of the main S2/L2 fabric they display. Samples considered in the calculations were garnet–chloritoid– paragonite micaschists west of the Cerro del Almirez ultramafic massif and north from the smaller bodies (Fig. 2a, see Fig. 7e, g, and h for photographs). Representative chemical analyses of minerals in different textural positions are shown in Table 2. The peak-metamorphic mineral assemblage consists of chloritoid and garnet porphyroblasts surrounded by the main S2 schistosity (Fig. 7e, g, and h). Large phengite grains (see Table 2 for composition) with a decussate texture obliterate the S2 foliation (Fig. 7b). Rutile, ilmenite, apatite, monazite, allanite, tourmaline, and zircons are accessory minerals. Garnet porphyroblasts show irregular inclusion-rich cores with quartz, phengite, minute chloritoid, chlorite, allanite, apatite, and rutile rimmed by ilmenite and tourmaline and idiomorphic to xenomorphic inclusion-poor coronas (Fig. 7h). Chloritoid porphyroblasts include the same minerals as garnet in addition to small idioblastic garnet grains or larger fragmented garnet porphyroblasts. A post-S2 static event is recorded by the following: recrystallization of the large phengite grains with decussate texture (Fig. 7b), the growth of idiomorphic coronas around rounded garnet cores (Fig. 7h), the formation of ilmenite rims around rutile grains, and the crystallization of large (up to 1 cm in size) and randomly oriented (post-S2) chloritoid idioblastic porphyroblasts (Fig. 7e and h). A late assemblage consisting of phengite, chlorite, and quartz can be found in shear zones, and replacing or filling open fractures in garnet and chloritoid porphyroblasts (Fig. 7e and h). Inclusion-rich cores of garnet porphyroblasts are almandine- and grossular-rich (XAlm = 0.69–0.80; XGrs = 0.11–0.25) and spessartineand pyrope-poor (XSps = 0.01–0.04; XPrp = 0.03–0.04). They show a weak typical prograde zonation (increasing Alm and Prp and decreasing Grs and Sps contents) towards the contact with the corona. Coronas are slightly richer in Prp and poorer in Grs than the cores (XAlm = 0.72–0.80; XGrs = 0.06–0.18; XSps = 0.01–0.02; XPrp = 0.08–0.16). The small garnet inclusions within chloritoid porphyroblasts display very homogeneous compositions (XAlm = 0.79–0.80; XGrs = 0.15–0.17; XSps = 0.01; XPrp = 0.06–0.07) within the range of those corresponding both to cores and coronas (Table 2, Fig. 7f). Chloritoid porphyroblasts show only limited zoning with an XMg variation ranging from 0.22 to 0.26 from core to rim (Table 2, Fig. 7i). Chloritoids included within garnets are Fe-richer, with XMg ranging from 0.11 to 0.17. Mn contents are always below 0.008 atoms per formula unit (apfu). According to the above-described textures, crystallization of the main mineral assemblage and the formation of the S2 foliation were simultaneous in the light-coloured schists in direct contact with the ultramafic bodies. We estimated the P–T conditions during the metamorphic evolution by calculating a phase diagram section (pseudosection) with PERPLEX (Connolly, 1990, 2009) for a representative bulk-rock composition, corresponding to one of the studied Tahal schist samples (SiO2 = 54.69 wt.%, Al2O3 = 24.83 wt.%, FeO = 7.47 wt.%, MgO = 1.43 wt.%, CaO = 0.99 wt.%, Na2O = 0.85 wt.%, K2O = 4.29 wt.%, and LOI = 3.01 wt.%). We used the thermodynamic database from Holland and

Powell (1998 upgraded in 2002) and appropriate solid solution models (see results in Fig. 8). The mineral assemblage garnet + chloritoid + phengite + paragonite + quartz (observed in the rock to be prekinematic to the main S2 foliation) is stable in a narrow P–T field at 1.3–1.9 GPa and 570–590 °C (Fig. 8a). These conditions are constrained by the stability fields of the minerals not observed in the rock: amphibole, at lower temperatures, and kyanite at higher temperature and higher pressure conditions. Syn-S2 metamorphism occurred at around 580 °C and 1.25 GPa, in the paragonite-absent field, in which phengite coexists with garnet + chloritoid + quartz and minor amounts of chlorite. The postkinematic path was characterized by the observed increasing amounts of chlorite in a wide field towards lower pressure (from 1.4 to 0.7 GPa) and temperature (from 580 to 460 °C). Further P–T constraints can be obtained from isopleths that show the predicted chemical variation of minerals in the pseudosection (Fig. 8b, c). The predicted composition of garnet (XAlm = 0.74–0.80, XPrp = 0.11–017, XGrs = 0.07–0.12; black isopleths in Fig. 8b) in the fields of interest (grey-shaded in Fig. 8b, c) fits that corresponding to the rims in the analysed garnets (Table 2). A similar agreement is found in coexisting paragonite and phengite (XPg = 0.81–0.90, XCel = 0.05–0.19; black isopleths in Fig. 8c) and also in chlorite (XMg = 0.56–0.60; grey isopleths in Fig. 8c). The predicted XMg range for chloritoid in the lower two pressure fields (0.21–0.29; grey isopleths in Fig. 8b) only partially fits the composition observed in the chloritoid porphyroblasts, probably reequilibrated with chlorite. In contrast, the predicted chloritoid composition in the peak metamorphic assemblage is richer in MgO (0.29–0.31) than the composition of the chloritoid inclusions found within garnet (Table 2). This is most probably a consequence of the ideal solid solution model of chloritoid (Ctd (HP), after White et al., 2000) available in the PERPLEX database. However, good agreement between the observed and the estimated compositions for all the other main minerals of the rock point to consistent results in the calculated pseudosection. In conclusion, a general PT path can be modelled for the studied light-coloured schists, in which peak-pressure conditions were reached above 1.3 GPa, and most probably, in accordance with the celadonite content of phengite, at around 1.9 GPa, and peak temperature was restricted to 570–590 °C (Fig. 9). This event was followed by the retrograde formation of chlorite, first during decompression and subsequently in a strong cooling and decompression event. The maximum temperature reached during the peak metamorphism, considerably lower than that deduced for ultramafic rocks, indicates that these two sequences were most probably juxtaposed after peak metamorphism was reached in both rock types (Fig. 9). 5. Discussion 5.1. Deformation of ultramafic rocks during progressive subduction The rheology of serpentinite is essential to understanding the behaviour of the slab-mantle wedge interface in a subduction zone. The study of natural serpentinites exhumed from ancient subduction zones (like those of the Cerro del Almirez massif) can help to determine how they behave in this setting. Three deformational fabrics in the ultramafic rocks of the Almirez massif and/or nearby bodies are relevant worldwide and they deserve a detailed discussion on their origin and significance in subduction settings: 1) The preservation of S1 foliation in serpentinites, which, based on field and textural relationships with successive metamorphic assemblages, can be related to the progressive burying of rocks until peak pressure is reached. This fabric can help to pinpoint the deformational evolution of serpentinites in a subduction zone. 2) The occurrence in these serpentinites of a set of joint fractures filled with olivine that, despite the lithostatic high-pressure conditions,

Table 2 Representative analyses of minerals from sample ALM01 from the light-coloured schists of the Mulhacén Unit. Minerals: Grt: garnet, Cld: chloritoid, Chl: chlorite, Ph: phengite, and Pg: paragonite. Textures: rim: rim of garnet porphyroblast, fr inc: fragment of garnet included in chloritoid, in Cld: small garnet included in chloritoid, core: core of garnet porphyroblast, porphy: chloritoid porphyroblasts, C: shear surface, and S2: S2 foliation. ALM01

ALM01

ALM01

ALM01

ALM01

ALM01

ALM1b

ALM1b

ALM01

ALM01

ALM01

ALM01

ALM01

ALM01

ALM1b

ALM01

ALM01

ALM01

ALM01

Mineral

Grt

Grt

Grt

Grt

Grt

Grt

Grt

Grt

Cld

Cld

Cld

Cld

Chl

Chl

Chl

Ph

Ph

Ph

Ph

ALM01 Pg

Texture

Rim

Rim

Fr inc

Rim

Fr inc

In Cld

Core

Core

Porphy

Porphy

Porphy

Porphy

C

C

C

S2

S2

C

C

C

Analysis

2

4

38

58

105

114

231

145

5

27

41

108

12

14

239

10

11

16

17

62

SiO2 TiO2 Al2O3 Cr2O3 FeO MgO MnO CaO Na2O K2O Total Si Al AlIV AlVI Ti Cr Fe2+ Fe3+ Mg Mn Ca Na K XMg XAlm XGrs XSps XPrp

36.93 0.04 20.82 0.00 36.81 0.51 1.55 4.48 – – 101.10 5.92 3.93 0.08 3.85 4.72 0.22 0.01 0.22 0.37 0.07 0.77 – – – 0.80 0.13 0.01 0.06

36.97 0.03 20.91 0.00 37.12 0.22 2.67 3.11 – – 101.00 5.91 3.94 0.10 3.84 4.71 0.25 0.00 0.25 0.64 0.03 0.53 – – – 0.80 0.09 0.01 0.11

37.32 0.04 21.07 0.02 34.54 0.25 3.75 3.59 – – 100.60 5.93 3.95 0.08 3.88 4.45 0.19 0.00 0.19 0.89 0.02 0.57 – – – 0.74 0.10 0.01 0.15

37.28 0.05 20.91 0.01 36.78 0.12 4.09 1.60 – – 100.80 5.94 3.93 0.06 3.87 4.65 0.18 0.01 0.22 0.98 0.02 0.30 – – – 0.79 0.05 0.00 0.16

37.18 0.07 20.84 0.00 35.18 0.41 1.46 5.98 – – 101.10 5.94 3.93 0.06 3.88 4.53 0.17 0.01 0.18 0.40 0.05 0.96 – – – 0.76 0.17 0.01 0.06

37.43 0.08 20.88 0.03 32.67 0.44 2.20 6.99 – – 100.70 5.96 3.92 0.04 3.87 4.19 0.15 0.01 0.15 0.52 0.06 1.19 – – – 0.70 0.20 0.01 0.09

37.21 0.13 20.66 0.04 31.23 1.54 0.88 8.79 – – 100.50 5.97 3.90 0.04 3.87 4.05 0.14 0.02 0.14 0.21 0.21 1.51 – – – 0.68 0.25 0.04 0.04

36.76 0.09 20.66 0.00 35.76 0.95 1.10 5.61 – – 100.90 5.92 0.32 0.09 0.23 0.01 0.00 4.58 0.23 0.23 0.15 0.97 – – – 0.77 0.16 0.03 0.04

24.81 0.01 40.87 – 22.49 4.05 0.06 0.00 0.00 0.01 92.30 2.03 3.00 – – 0.00 – 0.01 1.53 0.49 0.00 0.00 0.00 0.00 0.24 – – – –

24.64 0.00 40.75 – 22.93 3.87 0.10 0.02 0.01 0.01 92.33 2.02 3.00 – – 0.00 – 0.04 1.53 0.47 0.01 0.00 0.00 0.00 0.24 – – – –

24.49 0.01 40.60 – 22.75 3.90 0.09 0.00 0.00 0.01 91.85 2.01 3.00 – – 0.00 – 0.04 1.53 0.48 0.01 0.00 0.00 0.00 0.24 – – – –

24.53 0.00 40.70 – 22.84 3.83 0.11 0.02 0.00 0.02 92.05 2.01 3.00 – – 0.00 – 0.04 1.53 0.47 0.01 0.00 0.00 0.00 0.23 – – – –

25.08 0.06 23.11 – 24.47 13.89 0.05 0.02 0.01 0.04 86.73 2.64 2.87 1.35 1.52 0.01 – 2.05 0.11 2.18 0.00 0.00 0.00 0.01 0.52 – – – –

24.72 0.06 22.96 – 24.55 14.14 0.02 0.02 0.00 0.05 86.52 2.59 2.84 1.40 1.43 0.01 – 1.79 0.37 2.21 0.00 0.00 0.00 0.01 0.55 – – – –

24.51 0.06 22.43 – 26.28 13.98 0.05 0.00 0.01 0.03 87.35 2.57 2.78 1.42 1.36 0.01 – 1.96 0.35 2.19 0.00 0.00 0.00 0.00 0.53 – – – –

47.31 0.38 32.49 – 1.73 1.65 0.01 0.01 0.92 8.95 93.65 3.19 2.58 0.79 1.79 0.02 – 0.10 – 0.17 0.00 0.00 0.12 0.77 0.63 – – – –

47.78 0.43 32.18 – 1.49 1.65 0.00 0.00 0.96 9.12 93.84 3.21 2.55 0.77 1.79 0.02 – 0.08 – 0.17 0.00 0.00 0.13 0.78 0.66 – – – –

46.13 0.16 35.98 – 1.24 0.66 0.00 0.01 1.55 8.64 94.54 3.07 2.82 0.92 1.91 0.01 – 0.07 – 0.07 0.00 0.00 0.20 0.73 0.49 – – – –

46.32 0.28 35.13 – 1.34 0.85 0.00 0.00 1.42 8.74 94.27 3.10 2.77 0.89 1.88 0.01 – 0.08 – 0.09 0.00 0.00 0.18 0.75 0.53 – – – –

45.86 0.11 38.92 – 0.64 0.09 0.01 0.29 6.96 1.23 94.13 2.98 2.98 1.01 1.97 0.01 – 0.04 – 0.01 0.00 0.02 0.88 0.10 0.20 – – – –

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Sample

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point to the development of embrittlement processes due to fluid release during prograde metamorphism. This brittle deformation can help to reveal the relationship between fluid overpressure and the increase in permeability and fluid mobility in HP-metamorphic rocks. 3) The widespread development of grain-size reduction zones in Chlharzburgites, which attest to the pervasive release of fluids during antigorite breakdown, as well as to the prevalent low differential stresses during subduction-linked high-pressure metamorphism in these rocks. The GRSZs suggest that overpressured fluids and brittle deformations are not restricted to the serpentinites and can approach the lithostatic pressure favouring the formation of weak shear zones in the subduction slab.

a

5.1.1. Significance of the S1/L1 planar–linear fabric The asymmetric σ-type porphyroclasts and the S–C structures in the Atg-serpentinite indicate that the S1/L1 planar–linear fabric developed in a non-coaxial strain during the prograde metamorphic path compatible with depths of around 40–60 km in a subducting slab. However, L1 orientations show a high variability including the existence of two maxima at nearly 90° and the low value of the strength parameter C (C = 1.637). This strong dispersion in the trend and dips of the L1 stretching lineation cannot be explained by later folding as most of the Atg-serpentinites only crop out in the Cerro del Almirez mountain with a tabular shape (Figs. 2 and 4), suggesting that this dispersion within the S1 foliation surface was an original characteristic of the S1/L1 fabric. Similar fabrics containing a high variability in lineation trends within the foliation surface have been observed from several natural shear zones (Czeck and Hudleston, 2003; Díaz Azpiroz and Fernández, 2005; Sarkarinejad and Azizi, 2008), and are usually explained by transpressional deformation models (see Fernández and Díaz-Azpiroz, 2009; Jiang, 2007). We will return to the possibility of transpressional deformation when discussing the emplacement of the ultramafic massif over the metasedimentary sequence. The formation of the S1/L1 structures is likely related to HP/LT Alpine metamorphism during subduction and not before, as deduced from the relationship between S1/L1 structures and the crystallization of prograde minerals observed in the Atg-serpentinite. Textural features constrain the temperature of Atg-serpentinite deformation because brucite breakdown occurs at about 475 °C (López Sánchez-Vizcaíno et al., 2009), and tremolite formation in these rocks occurs at 600–630 °C and 1.6–1.9 GPa (Padrón-Navarta et al., 2012). This is consistent with the overprint of the S1/L1 fabric and the WNW– ESE open folds by the antigorite-out isograd, as described in Section 3.1, which took place at around 680 °C at the same pressure (Padrón-Navarta et al., 2011). Taking into account the above metamorphic reactions, we can conclude that ductile deformation in Atg-serpentinite can be accommodated at around medium-grade conditions (475–630 °C). The non-coaxial deformation is in accordance with the model from Angiboust et al. (2012) of shearing on top of a subducting slab facilitated by mechanical weakening resulting from the serpentinization of ultramafic rocks. This is in spite of the differences between the lithological sequence considered in this model and that of the Cerro del Almirez massif. 5.1.2. Veins and hydrofracturing in the Atg-serpentinite After the generation of the S1/L1 fabric, static annealing followed at circa 680 °C and 1.6–1.9 GPa (Padrón-Navarta et al., 2012). However, brittle veins appear crosscutting the S1 foliation. As previously stated,

b

c

Fig. 8. a) P–T pseudosection (system SiO2–Al2O3–FeO–MgO–CaO–Na2O–K2O) computed with PERPLEX (Connolly, 2009) for a Tahal schist sample of the Mulhacén Unit in direct contact with the ultramafic rocks of the Cerro del Almirez área. White, light grey, intermediate grey, and dark grey fields represent stability variances of two, three, four, and five, respectively. The high-pressure garnet + chloritoid + phengite + paragonite + quartz ± H2O assemblage (in red) is prekinematic to the main S2 foliation. Syn-S2 metamorphism occurred at around 580 °C and 1.25 GPa in the field without paragonite, and the postkinematic path is characterized by increasing amounts of chlorite (not shown in this figure) at lower P and T. Numbered fields correspond to the following assemblages: 1: Cld Amp Ms Grt Ky Qz, 2: Chl Cld Amp Amp Ms Grt Ky Qz, 3: Bt Cld Amp Amp Ms Grt Ky Qz, 4: Bt St Cld Amp Amp Ms Grt Qz, 5: Chl St Cld Amp Ms Grt Qz, 6: Chl Cld Amp Ms Grt Ky Qz, 7: Bt St Ms Grt Qz H2O, 8: Bt St Fsp Ms Grt Qz H2O, 9: Bt Fsp Ms Grt Sil Qz H2O, 10: Bt melt Fsp Ms Grt Sil Qz H2O, 11: melt Crd Fsp Fsp Grt Sil Qz, 12: melt Crd Fsp Fsp Grt Sil, 13: melt Fsp Fsp Grt Sil, 14: melt Fsp Fsp Grt Sil Qz, 15: melt Fsp Grt Sil Qz, 16: melt Fsp Grt Ky Qz. b) Isopleths displaying the calculated compositional variation of garnet for the almandine (XAlm), pyrope (XPrp), and grossular (XGrs) components and for XMg in chloritoid. c) Isopleths displaying the calculated compositional variation of white micas for the celadonite (XCel), and paragonite (XPg) components and for XMg in chlorite. The shaded field corresponds to the stability conditions of the mineral assemblages observed in the rock. Abbreviations for mineral names after Whitney and Evans (2010).

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2.00

Atg-serpentinites S1/L1fabric & jointing 18-15 Ma GSRZ

0.50 400

Atg Chl Ol Cpx H2O

500

Ol T r H2 O Atg-o ut

P(GPa)

F3 folds

Chl Ol Tr Tlc H2O

600

%volume H2O 50

40

30

rgites

1.00

Ol Opx Chl Cpx H2O

arzbu

Tahal schists S2/L2 fabric

1.25

0.75

Chl-h

1.50

Atg Ol Chl Brc Cpx

Atg Chl

1.75

20

Ol Opx Chl Tr H2O

10

Ol Chl Tr Ath H2O

700

800

Fig. 9. Estimated P–T path for the Tahal schists (black line; see Fig. 8) compared with the P–T path from the ultramafic rocks in the Cerro del Almirez massif, where the green line marks the prograde P–T path for the Atg-serpentinites estimated by López SánchezVizcaíno et al. (2009) and Padrón-Navarta et al. (2012), and the orange line marks the retrograde P–T path for the Chl-harzburgites estimated by Padrón-Navarta et al. (2010a). The yellow line marks a likely retrograde trajectory for the ultramafic rocks that joined the Tahal schist path at the development conditions of the S2/L2 fabric (580 °C and 1.25 GPa, see Fig. 8). The P–T paths are superposed over the modal variations of H2O (in vol.%) released during dehydration of the ultramafic rocks and calculated for the composition of a representative serpentinite from the study area on both sides of the antigorite-out isograd (López Sánchez-Vizcaíno et al., 2009).

these veins are filled with products of prograde dehydration reactions from both brucite breakdown (López Sánchez-Vizcaíno et al., 2009) and partial antigorite dehydration reactions before the final breakdown of this mineral. The afore-described features of the vein structure (such as the olivine crystals with a fibrous or elongated habit oriented normal to the walls, angular irregularities in sharp vein walls that match across the vein, and curved vein tips (Fig. 6b–d)) are among the evidence that Healy et al. (2009) found for the brittle fracture origin of the dehydration veins in the Erro-Tobbio serpentinites (Alps). The structural association for these serpentines is also the same as those described in the Cerro del Almirez, with an older schistosity in the serpentinites crosscut by later brittle veins filled by metamorphic olivine ± Ti-clinohumite (Auzende et al., 2006; Healy et al., 2009; Scambelluri et al., 1995), and point to a similar origin. Moreover, hybrid joints with evidence of both opening normal to and shearing parallel to the joint walls are observed in the Atg-serpentinite (Fig. 6d and f). These brittle structures indicate the coexistence of both brittle and ductile deformations during the formation of S1. This contradictory behaviour is also documented in other serpentinite massifs (Neufeld et al., 2003; Rebay et al., 2012) and agrees with experimental studies on serpentinites below the PT conditions of antigorite breakdown (Escartin et al., 1997). A mechanism for the formation of such brittle structures and their relationships with the S1 foliation begins with the release of the fluid phase in the impermeable serpentinites, which in turn increases the pore pressure to the point of brittle rupture and also increases permeability, thus allowing the fluids to escape (e.g., Etheridge et al., 1984; Gudmundsson, 2011; Healy et al., 2009). The subsequent growth of olivine ± Ti-clinohumite seals the fractures and induces a renewed decrease in permeability, again allowing the fluid phase to increase pore pressure (e.g., Nollet et al., 2005). Fig. 9 depicts the estimated metamorphic P–T trajectories of the metasediments (dashed black line) and ultramafic rocks (green line for serpentinites and orange for harzburgites; Padrón-Navarta et al., 2010a, 2012) superposed over the modal variations of H2O (in vol.%) for a representative serpentinite from the study area (López Sánchez-Vizcaíno et al., 2009). Fig. 9 also

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shows how, during the prograde metamorphism of serpentinites and the generation of the planar–linear fabric (S1/L1), the predicted volume of released H2O ranges between 0 and 5% at low temperatures and rises to 5–10% when approaching the Atg-out isograd, allowing the generation of high fluid pressures. 5.1.3. GSRZ and hydrofracturing in the Chl-harzburgite Brittle structures reported here provide important information about the deformational regime that prevails beyond the breakdown of hydrated ultramafic rocks. Grain-size reduction zones (GSRZ) (see Section 3.2) are common in the prograde Chl-harzburgite. Despite their apparent ductile appearance (apparent mylonites), they record an episode of brittle deformation at high pressure and temperature (Padrón-Navarta et al., 2010a). GSRZs are systematically oriented and distributed in a girdle with two major sets of conjugate surfaces at circa 40° to each other (Fig. 3e and f), without any evidence of relative displacement (Padrón-Navarta et al., 2010a). Systematically oriented tension joints (like the GSRZs) are documented in sedimentary rocks from the uppermost levels of the crust and in cooled metamorphic and igneous rocks undergoing exhumation (e.g., Mandl, 2005). However, hydrofractures can form in the whole crust or even in the upper mantle (e.g., Healy et al., 2009) if the fluid pressure is high enough to produce the rupture of the rock in extension (e.g., Gudmundsson, 2011; Healy et al., 2009). Fig. 9 shows the significant increase in the volume of fluids released during the dehydration process, up to more than 25% of the rock volume, enough to induce hydrofracturing. Therefore, during the progression of the dehydration front, the high amount of fluid released during the serpentinite dehydration to produce olivine + pyroxene + chlorite + fluids (Padrón-Navarta et al., 2010a) accounts for the overpressured fluids and the generation of the GSRZs in the previously formed Chl-harzburgites. Hydrofractures in the upper mantle related with the antigorite dehydration in subduction zones have been proposed by Nishiyama (1989). Miller et al. (2003) also propose a model where the fluid pressures generated by serpentinite dehydration (among other minerals) are sufficient to induce hydrofracturing, which significantly influences the porosity–permeability structure, a fact they linked to seismicity in subduction zones and the observed post-seismic evolution of wave velocities (Vp/Vs) in major earthquakes. The geometry of these GSRZ consists of two sets of conjugate fractures with low (b60°) dihedral angles. This geometry is the same as that of conjugate hybrid joints, where the σ1 bisects this dihedral angle, σ3 bisects the supplementary angle, and σ2 is located at the intersection of the two sets of fractures (Muehlberger, 1961). The presentday orientation of the GSRZ conjugate sets indicates a roughly subhorizontal orientation of σ1 with a N–S trend, a sub-vertical orientation of σ3, and σ2 sub-parallel to the pole of the girdle distribution: a line with a N110°E trend and 30° dip towards the ESE (Fig. 3e and f). Those orientations may not coincide with their orientation when the structures formed as the rocks may have undergone some rotation during exhumation. Therefore, the orientations of the main stress axes might agree with a compressional setting for the hydrofracturing event. Similar σ1 stress axes with low dip angles are also predicted inside a subducting lithospheric slab (e.g., Faccenda et al., 2009; Gerya et al., 2008). Whatever the original orientation of the GSRZ conjugate sets, the formation of conjugate hybrid joints occurs when differential stresses are small (Muehlberger, 1961), suggesting that the effective stresses were also small during the increasing pressure conditions of both Atgserpentinites and Chl-harzburgites ultramafic rocks. Furthermore, Behr and Platt (2013) estimate low differential stress magnitudes (≤ 10 MPa) in the Nevado–Filábride Complex metasediments in the early stages of exhumation. Low differential stresses during highpressure metamorphism within a subducting slab were also recorded in metapelites from the Eclogite Zone cropping out in the Tauern Window (Alps) (Stöckhert et al., 1997) and in rocks from the Franciscan

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Complex (Grigull et al., 2012). In fact, Stöckhert (2002) and Wassmann and Stöckhert (2013) predicted that the HP metamorphic rocks have probably never experienced differential stresses exceeding a few to tens of megapascals during their entire subduction and exhumation history. This would imply that weak shear zones might occur between the essentially undeformed rock bodies of the subducting slab (Grigull et al., 2012; Stöckhert, 2002). 5.2. Emplacement of the ultramafic lens over the metasedimentary sequence In the study area, the emplacement of the ultramafic rock lenses over the schists was due to a major shear zone (Figs. 2 and 4) related with the S2/L2 fabric in the metasediments, which can be correlated to the above-mentioned weak shear zones occurring around essentially undeformed rock bodies. In fact, close to the boundary with the ultramafic rocks, metasediments display the penetrative S2 foliation parallel to the orientation of the major shear zone, as well as widespread S–C fabrics and other similarly oriented shear-related structures (Fig. 4). These shear structures cannot be observed within the Chl-harzburgite bodies that, according to the rheology of the olivine (e.g., Korenaga and Karato, 2008), were essentially brittle at the PT conditions estimated for D2 deformation. The common orientation of all these structures and their shear-related origin indicate that this major shear zone formed during the D2 deformational event in a non-coaxial regime with a top-to-the-west sense of movement. Peak metamorphic conditions deduced in the light-coloured schists are at lower temperatures (570–590 °C and 1.3–1.9 GPa) than in the ultramafic rocks (680–710 °C and 1.6–1.9 GPa, Padrón-Navarta et al., 2010a, 2012) (Fig. 9), suggesting that the two blocks were originally at a similar depth but at different temperatures within the subducting slab. These P-T conditions would be more consistent with the typical thermal structure of a hot subduction zone than with that of a cold subduction (e.g. Hacker et al., 2003), in agreement with detailed petrological data obtained from the Cerro del Almirez ultramafic rocks (López Sánchez-Vizcaíno et al., 2005, 2009). The fact that both the ultramafic and the metapelite blocks, with similar pressure conditions, were juxtaposed at the end of the D2 deformational event suggests that the shear zone had a strong strike–slip component, which in turn points to a transpressional deformation regime for the subducting slab. This is in agreement with the numerous geodynamic models proposed for the evolution of the Betic–Rif orogen, among others: subduction roll-back, slab break-off, or lithospheric delamination (see Platt et al., 2013 for a review). All these models propose the advance of the continental crust of the Alborán Domain towards the west between the European and African plates (e.g., Balanyá et al., 2007; Booth-Rea et al., 2005; Chalouan et al., 2008; Platt et al., 2003, 2013). This motion implied an oblique collision in the northern branch of the Betic–Rif orogen, generating a right-lateral transpressional regime affecting the South Iberian palaeomargin. Transpressional deformation could also explain the high variation in the trends of the L2 stretching lineation through the entire sequence of the Nevado–Filábride Complex from NW–SE in the eastern outcrops to WSW–ENE in the westernmost ones (see Jabaloy et al., 1993, their Fig. 2). 5.3. Subduction and exhumation of the upper Nevado–Filábride rocks and regional implications A series of deformational structures and fabrics preserved in the wall-rock metasediments allow us to deduce the emplacement mechanism of the ultramafic rocks in the Nevado–Filábride lithological sequence and the subsequent exhumation of the entire complex in the context of the Betic–Rif Chain evolution during the Alpine orogeny. PT conditions and geothermal gradients recorded in the sequence indicate that the most likely setting for the HP/LT Alpine metamorphism and deformations undergone by the Nevado–Filábride Complex

was the lower plate of a subduction zone while the upper plate corresponded to the Alborán Domain. Several additional data support this interpretation: the Nevado–Filábride metasedimentary sequence represents a Paleozoic series of the Iberian Massif passive margin, as indicated by Early Permian leucogranitoid intrusions in the metasediments (Gómez-Pugnaire et al., 2004, 2012). Zircon ages dating the HP event from gneisses, eclogites, and clinopyroxenites indicate that this palaeomargin was subducted beneath the Alpujárride and Maláguide complexes during the formation of the Gibraltar Arc at 18 to 15 Ma ago (Behr and Platt, 2012; Gómez-Pugnaire et al., 2004, 2012; López Sánchez-Vizcaíno et al., 2001; Platt et al., 2006). The original location of the Cerro del Almirez ultramafics prior to their emplacement over the Mulhacén metasedimentary sequence is not well constrained. These ultramafic rocks derived from mantle peridotites that underwent up to 20% partial melting in the spinel stability field (Marchesi et al., 2013). According to Alt et al. (2012) and Marchesi et al. (2013), these melt residues were originally exhumed to the seafloor in a continental passive margin or oceanic basin and were hydrated at different temperatures in a fluid-dominated system. Both the exhumation in an oceanic seafloor and the rodingitization of basic dykes, now occurring as rodingite boudins both in serpentinites and Chl-harzburgites, exclude the possibility that those ultramafic rocks formed part of the mantle wedge over the subduction slab, in contrast to the proposal of the Ronda Peridotite massif as a fragment of the subcontinental lithospheric mantle below the Alborán domain (e.g. Garrido et al., 2011). As reported in this work, each rock type in the lithological sequence of the Cerro del Almirez area accounts for different tectonometamorphic conditions and very likely for different stages in the complex evolution of a subduction setting and the subsequent exhumation process. In the study area, the serpentinites preserve the older prograde structures whereas the metasediments record the structures responsible for the exhumation. This situation is in clear contrast with other orogenic belts where extensive occurrences of serpentinites are thought to be responsible for the exhumation of large HP units (e.g., Angiboust et al., 2012; Guillot et al., 2009). When comparing the trend and sense of shear of all the non-coaxial deformational stages, we can see how those that formed first during subduction (D1) and those that formed later during exhumation in the subduction channel (D2) have the same kinematics with a top-to-thewest sense of movement and most likely represent a right-lateral transpressional deformation. Two opposite models are used to explain the exhumation of rocks deformed between two rigid plates: the subduction channel model (Shreve and Cloos, 1986), where shearing has the same vorticity in both rigid walls, and the orogenic channel model (e.g., Grujic et al., 2002), where the shearing has opposite vorticities in the walls limiting the rocks to be exhumed. The structures of the Cerro del Almirez massif are more in accord with a subduction channel geometry as there is no evidence of opposite vorticities between the prograde- and the retrograde-generated structures. Exhumation was accommodated by non-coaxial deformations superposing different tectonic units, suggesting that the movement of the rocks was accommodated by decoupling at shear zones and the superposition of thin tectonic units with slightly different P–T conditions. However, this mechanism did not produce the characteristic block-in-matrix structures of the chaotic HP–LT mélanges (e.g., Grigull et al., 2012) but instead a thin-layered structure of amalgamated old tectonic units. Guillot et al. (2009) proposed that the decoupling of the upper part of continental subducting lithosphere occurs at depths of 90–140 km, but our work indicates that decoupling can also occur at lower depths of 36–70 km. Only when the rocks of the Nevado–Filábride Complex reached upper crustal conditions at 12–9 Ma was exhumation accommodated by brittle extensional detachments as described extensively by several authors (e.g., Augier et al., 2013; Do Couto et al., 2014; Galindo-Zaldívar et al., 1989; Jabaloy et al., 1993).

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6. Conclusions The structural analysis in the Cerro del Almirez area provides information about deformation mechanisms during the subduction of a continental sequence that includes relatively small slivers of ultramafic rocks. Evidence of ductile mesostructures during the prograde path are preserved only within the ultramafic lithologies and include noncoaxial deformation in the Atg-serpentinites. In these rocks, the simultaneous release of fluids during the metamorphic dehydration of brucite and the partial breakdown of antigorite produced embrittlement conditions with the generation of olivine ± Ti-clinohumite veins and hybrid joints. At higher temperatures corresponding to the antigorite-out isograd, rigid Chl-harzburgites formed. The associated fluid released during the antigorite breakdown triggered hydrofracturing of Chlharburgites in a compressional setting with low differential stresses and a probably sub-horizontal σ1. Ductile exhumation of the rocks is mainly preserved within the metapelites and marbles involving a recurrent sequence of noncoaxial deformations and folding. The first fabric to develop during exhumation is the S2/L2 planar–linear fabric formed within shear zones dominated by a non-coaxial regime with a top-to-the-west sense of movement in a compressive regime at 1.3–1.9 GPa and 570–590 °C. This fabric is associated with the final configuration of the subducting continental sequence consisting of superposed thin tectonic units including both a metasedimentary sequence and ultramafic rock lenses recording different metamorphic conditions and, most likely, a righthanded transpressional deformation. The structures described in this work indicate that decoupling and folding worked in the upper part of a continental subducting slab at depths of 36–70 km with a geometry similar to that of subduction channels. Decoupling of the upper part of the continental subducting slab by shearing and folding is a likely mechanism to produce the exhumation of thin tectonic units. As documented in this work, non-coaxial deformation stages related to subduction and exhumation can have similar kinematics. In the study area, the serpentinites preserve the old prograde structures whereas the metasediments record the structures responsible for the exhumation. This situation is in clear contrast with other orogenic belts dominated by large ophiolite sequences where extensive occurrences of serpentinites are thought to be responsible for the exhumation of large HP units. Acknowledgements This work has been supported by the Spanish MICINN research grants CGL2011-24101, CGL2011-29920, and CGL2012-32067 (financed by the European Regional Development's funds), and by the Junta de Andalucía research groups RNM-148, RNM-145, and RNM131 and grant Proyecto de Excelencia-2009-RNM-4495. Alexander Robinson and an anonymous reviewer are thanked for helpful suggestions that improved the manuscript. Christine Laurin is thanked for revising the English. Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.tecto.2014.12.022. References Agard, P., Yamato, P., Jolivet, L., Burov, E., 2009. Exhumation of oceanic blueschists and eclogites in subduction zones: timing and mechanisms. Earth Sci. Rev. 92, 53–79. http://dx.doi.org/10.1016/j.earscirev.2008.11.002. Agard, P., Augier, R., Monié, R., 2011. Shear band formation and strain localization on a regional scale: evidence from anisotropic rocks below a major detachment (Betic Cordilleras, Spain). J. Struct. Geol. 33, 114–131. http://dx.doi.org/10.1016/j.jsg.2010. 11.011. Alt, J.C., Garrido, C.J., Shanks III, W.C., Turchyn, A., Padrón-Navarta, J.A., López SánchezVizcaíno, V., Gómez-Pugnaire, M.T., Marchesi, C., 2012. Recycling of water, carbon,

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