Using 3D kinematic models in subduction channels. The case of the Chañaral tectonic mélange, Coastal Cordillera, northern Chile

Gondwana Research 74 (2019) 251–270

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Using 3D kinematic models in subduction channels. The case of the Chañaral tectonic mélange, Coastal Cordillera, northern Chile Paulina Fuentes a, Carlos Fernández b, Juan Díaz-Alvarado a,⁎, Manuel Díaz-Azpiroz c a b c

Departamento de Geología, Universidad de Atacama, Copayapu 485, Copiapó, Chile Departamento de Ciencias de la Tierra, Universidad de Huelva, 20171 Huelva, Spain Departamento de Sistemas Físicos, Químicos y Naturales, Universidad Pablo de Olavide, 41013 Seville, Spain

a r t i c l e

i n f o

Article history: Received 11 September 2018 Received in revised form 12 December 2018 Accepted 15 December 2018 Available online 3 February 2019 Keywords: Tectonic mélange Subduction channel Kinematic models Chañaral mélange Northern Chile

a b s t r a c t The Chañaral tectonic mélange (northern Chile) is a local unit within the late Paleozoic accretionary complex formed at the southwestern margin of Gondwana. The structural characteristics of the studied mélange were mostly developed during a first deformation phase (D1) and include a block-in-matrix fabric, lineations (L1) and foliations (S1), tight to intrafoliar folds, S-C and S-C-C′ composite planar fabrics, and a conspicuous spatial separation of domains with predominantly linear and linear-planar fabrics. Folding during a second stage (D2) modified the orientation of the previous fabrics and structures. The eastern boundary of the Chañaral mélange is N-S to NNW-SSE oriented, moderately dipping to the east. Its western boundary is not exposed. The plane showing the maximum structural asymmetry (the vorticity normal section) is ENE-WSW directed, and subvertical. Kinematic criteria consistently reveal top-to-the-WSW displacement. A kinematic model of triclinic transpression with inclined extrusion has been applied to evaluate the D1 structural features of the Chañaral mélange. The pitch of the simple shear direction on the deformation plane ranged from 60°N to 90°. The pitch of the estimated extrusion direction was of 30°-40°S. The coaxial component was clearly constrictional (logarithmic K value of 2 to 5). The vorticity number has not been constrained by the model, but its spatial variation can explain the domainal distribution of the fabrics in the mélange. The simulated particle paths show the predominance of material displacement parallel to the margin, with low to moderate down-dip displacements, which is in accordance with the low-pressure metamorphic assemblages found in the mélange. The convergence direction between the blocks separated by the mélange unit was N50°-60°E. Kinematic blocking of the Chañaral mélange, probably related to the accretion of an oceanic volcanic domain, allowed the D2 folding of the previous structures, a process that, at least initially, proceeded without a change in the convergence direction. © 2019 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction Mélanges and broken formations have been related to diverse tectonic settings and structural levels as the products of intense deformational processes in ancient and recent convergent margins and orogens (e.g., Raymond, 1984; Cowan, 1985; Festa et al., 2010; Kimura et al., 2012). The term has received quite different meanings (e.g., wildflysch, argille scagliose, olistostromes, megabreccias and agglomerates) depending on the focus of the studies (lithology, structure, tectonic, and sedimentary interpretation) or the processes involved in the formation of the block-in-matrix textures that characterize most of the mélange-like units (e.g., Cloos, 1982; Suzuki, 1986; Raymond et al., 1989; Camerlenghi and Pini, 2009; Festa et al., 2010). However, recently, the term mélange has been referred to mainly descriptive ⁎ Corresponding author. E-mail address: [email protected] (J. Díaz-Alvarado).

aspects providing, in turn, a classification of types and subtypes that includes a great diversity of tectonic contexts and leads to more precise interpretations (Festa et al., 2012). In accretionary complexes, mélange are key formations to understand the processes that govern the subductive margins, as the convergence direction, the down- and up-ward paths of subducted material, the degree of consolidation of trench-fill sediments or seismicity (e.g. Cloos, 1982; Niwa, 2006; Meneghini et al., 2009; Kimura et al., 2007, 2012; Krohe, 2017). According to their origins, Festa et al. (2010, 2012) mainly identified two mélange subtypes related to subduction in accretionary prisms. Mass-transport deposits at the frontal accretionary prism (Abbate et al., 1970; Pini, 1999; Festa et al., 2010, 2012) and tectonic mélanges and broken formations that show structural and lithological evidence of various degrees of mixing and deformation depending on, among others, the nature and rheology of the involved materials, the superposition of different tectonic stages and the transient or discontinuous participation of internal processes associated

https://doi.org/10.1016/j.gr.2018.12.009 1342-937X/© 2019 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

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with the plate boundaries (Lash, 1987; Cousineau, 1998; Pini, 1999; Yamamoto et al., 2000, 2009; Ujiie, 2002; Niwa, 2006; Grigull et al., 2012). Many studies have focused on detailed structural analyses of tectonic mélanges related to the underplating of sedimentary sequences in subductive margins (e.g., Cowan and Brandon, 1994; Onishi and Kimura, 1995; Kusky and Bradley, 1999; Kano and Konishi, 2001; Shi et al., 2013; Escuder-Viruete et al., 2013; Escuder-Viruete and Baumgartner, 2014; Zheng and Zhao, 2017). At the meso-scale, most of the structures described in the literature are highly similar, mainly pinch-and-swell structures and boudinage of the original planar bedding, isoclinal folds, sheath folds, S-C-C′ structures, slump folds, slump balls, extensional duplexes and thrust systems, veins and joint systems, hackle fringes and plumose structures (Hsü, 1968; Raymond, 1984; Cowan, 1985; Taira et al., 1992; Pini, 1999; Vannucchi and Bettelli, 2002; Ujiie, 2002; Bettelli and Vannucchi, 2003; Yamamoto et al., 2007, 2009; Ohsumi and Ogawa, 2008; Festa et al., 2010). As a seismogenic zone, the consideration of tectonic mélanges as fault rocks is an updated and crucial subject (Kimura et al., 2007, 2012; Kitamura and Kimura, 2012). Meanwhile, the variables that determine the general behavior of the subduction channel and their relevance in the subduction deformational processes are still to understand, although significant efforts have been carried out to assess the influence of internal parameters of the subductive margins in the tectonic mélanges, such as the angular relations between the wedge orientation and the plates convergence vector, the amount of deformation occurred during the mélange formation, or the observed variations between sediment-starved and sediment-filled channels (Cloos, 1982; Meneghini et al., 2009). However, as a transient and evolutionary process, the distribution of lithologies, the degree of consolidation of the sediments, the metamorphic grade, the strain rate and partitioning, the fluidization, the nature of the back-stop and other characteristics of the process may change at some time during mélange evolution depending on time and location within the subduction channel (e.g., Shreve and Cloos, 1986; Cloos and Shreve, 1988a, 1988b; Niwa, 2006; Festa et al., 2012; Kimura et al., 2012; Krohe, 2017). One of the main features to consider in the structural study of the tectonic mélange formations when we evaluate them as deformation zones is the ubiquitous oblique convergence in the subductive plate limits (Philippon and Corti, 2016). Although obliquity typically leads to 3D deformation (e.g., Jiang et al., 2001; Díaz-Azpiroz et al., 2016, 2018), very few studies have delved into the 3D distribution of the mélange processes (Cloos, 1982; Niwa, 2006; Plunder et al., 2018), and most of them present 2D models orthogonal to the trench (but not necessarily parallel to the convergence direction) or they consider an orthogonal convergence. To fill this gap, the model presented in this work takes into account a deformed rock volume in 3D, under different starting conditions of the main kinematic parameters controlling the flow. To adequately analyze the 3D (or, at least, 2.5D) kinematics of complex deformation zones, as is presumably the case of most tectonic mélanges, it is necessary to use modern analytical approaches, including analytical mechanical modeling (e.g., Fitch, 1972; McCaffrey, 1992; Enlow and Koons, 1998; Haq and Davis, 2010), analogue models (e.g., Schreurs and Colletta, 1998; Casas et al., 2001; McClay et al., 2004; Leever et al., 2011), numerical models (e.g., Braun, 1993; McCaffrey et al., 2000), and kinematic modeling of transpression and transtension (e.g., Fossen and Tikoff, 1998; Jiang and Williams, 1998; Jiang et al., 2001; Fernández and Díaz-Azpiroz, 2009). The latter models have decisively contributed to a better understanding of the complex deformational settings predominating in oblique convergent settings (Díaz-Azpiroz et al., 2016). This article presents a structural analysis of the Chañaral tectonic mélange in northern Chile, part of the late Paleozoic accretionary complex in the southwestern margin of Gondwana. The study has been carried out by evaluating the kinematic model of triclinic transpression with inclined extrusion described by Fernández and Díaz-Azpiroz

(2009), which has been successfully applied in different natural cases (e.g., Fernández et al., 2013; Díaz-Azpiroz et al., 2014; Alonso-Henar et al., 2018). The orientation, arrangement and vergence of the tectonic structures and fabrics measured in the mélange unit at Chañaral and Obispito localities conform a robust dataset for the application of the model, which allows us to obtain a three-dimensional perspective of the complex relationship between the oblique plate motion and the final structural configuration of the tectonic mélanges. 2. Geological setting 2.1. Tectonic overview The Chañaral tectonic mélange belongs to the western domain of the Las Tórtolas Formation (Ulricksen, 1979; Bell, 1982), also known as the Chañaral Epimetamorphic Complex (Godoy and Lara, 1998), one of the metamorphic complexes that are discontinuously exposed along the north and central Chilean coast (≈26°-39°S) and are associated with the accretionary processes occurred in the southwestern margin of Gondwana during the late Paleozoic (Fig. 1a, b), which was previously marked by the accretion of terranes (Charrier et al., 2007; Hervé et al., 2007, 2013; Heredia et al., 2018). A subductive setting was established along the southwestern margin of Gondwana between 345 and 335 Ma (Bahlburg et al., 2009; Heredia et al., 2016) while the first metamorphism and magmatism related to the convergent margin have been estimated around 320 Ma (e.g. Hervé et al., 1984, 1988; Bahlburg and Hervé, 1997; Willner et al., 2004, 2005). These late Paleozoic accretionary complexes are mostly comprised by thick turbiditic sequences deposited during the previous (early Carboniferous) passive margin stage and the trench-fill sediments that fed the accretionary system (e.g., Bahlburg et al., 2009; Hervé et al., 2013; Pankhurst et al., 2016). In addition, oceanic basaltic rocks appear as tectonic slices and deformed blocks in the metaturbiditic matrix with different metamorphic grades (e.g. Hervé, 1988; Willner, 2005; Hyppolito et al., 2014; Fuentes et al., 2016). The coastal metamorphic complexes are conformed by two main units: 1) the so-called Western Series present highly deformed highgrade (HP-LT) metamorphic rocks in south-central Chile and lowgrade mélange facies and volcanic rocks to the north (Las Tórtolas Formation, Fig. 1c). Metamorphic peak conditions in these HP-LT western units reached up to 11 kbar and 700 °C (e.g. Willner, 2005; Creixell et al., 2016). 2) Low grade metasedimentary units are located to the east (ca. 5 kbar, 250–300 °C, Willner et al., 2001, 2012), called Eastern Series (Fig. 1b) (Aguirre et al., 1972; Hervé, 1988; Willner, 2005). In the Accretionary Complex of Central Chile (34°-39°S) (Fig. 1b), the low-grade metasedimentary units were overprinted by hightemperature conditions during the intrusion of the subduction-related batholith (e.g. Willner, 2005; Hyppolito et al., 2015). Western and Eastern series are interpreted as the basal and the frontal accretionary formations respectively of an east- to northeast-directed subductive margin, active between the middle Carboniferous and the earlymiddle Permian (e.g., Willner, 2005; Willner et al., 2005, 2012; GarcíaSansegundo et al., 2014; Hyppolito et al., 2014, 2015; Fuentes et al., 2016, 2017; Creixell et al., 2016; Heredia et al., 2016). This period of activity of the accretionary processes along the margin between 26° and 39°S is characterized by a quite similar tectonic evolution of the entire convergent margin, i.e. homogeneous activity of the processes of accretion, metamorphism and magmatism (Gondwanan cycle, e.g. Ramos, 1988; Charrier et al., 2007), giving place to a first deformational stage. During early-middle Permian times, the final exhumation of the basal units (Western Series), i.e. the second deformational stage, constituted the accretionary complexes with significant N-S differences in the contact relations between the HP and LP metasedimentary units and accreted oceanic rocks (e.g. Creixell et al., 2016; García-Sansegundo et al., 2014; Hyppolito et al., 2015; Fuentes et al., 2016; Heredia et al., 2018). The subsequent middle

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Fig. 1. (a) Map of Gondwana and location of the study area (modified from Lawver and Scotese, 1987). (b) Distribution of the accretionary complex fragments along the Chilean coast and the late Paleozoic to Upper Triassic magmatism associated with the convergent margin. The best studied fragments of this ancient accretionary setting are, from north to south: the Las Tórtolas Formation, also known as the Chañaral Epimetamorphic Complex (26°-29°S), the Punta de Choros Metamorphic Complex (28°30′-29°30′S), the Choapa Metamorphic Complex (31°30′S) and the Coastal Accretionary Complex of Central Chile (34°00′-38°30′°S) (e.g., Fuentes et al., 2017; Creixell et al., 2016; García-Sansegundo et al., 2014; Willner, 2005). (c) Geological map of the study area. The detailed structural studies were carried out in Chañaral and Obispito areas.

Permian-Upper Triassic phase (pre-Andean cycle) was determined by the margin segmentation according to the differences in the tectonic activity and the nature, volume and location of the magmatism (e.g., DíazAlvarado et al., 2019). The pre-Andean tectonic scenario is characterized by crustal extension, dextral block rotations (e.g. Kleiman and Japas, 2009), N-S dextral shear zones (Kato and Godoy, 2015) and the sinistral displacement along the NW-SE Lanalhue Fault Zone to the south (≈39°S) (Glodny et al., 2008). In general, the accretionary processes were absent during this latter period and, mostly in northern Chile, the previous metamorphic complex was affected by high-silica intrusives with mantle isotopic signatures (Hervé et al., 2014; del Rey et al., 2016). The basal accretion of basaltic oceanic rocks (metabasites) was active along the margin during most of the late Carboniferous period (e.g. Willner, 2005; Willner et al., 2008; Hyppolito et al., 2014, 2015; Creixell et al., 2016), whereas the final cessation of the accretionary processes (early to middle Permian) coincided with the arrival of topographic highs, oceanic plateaus or seamounts to the margin, which are included as frontally accreted and obducted basaltic units in the Eastern Series (Heredia et al., 2016; Díaz-Alvarado et al., 2019). The collision of this oceanic ridge or seamounts with the margin may have promoted

the important changes occurred in the western margin of Gondwana during the pre-Andean cycle (García-Sansegundo et al., 2014; Hyppolito et al., 2014; Heredia et al., 2016; Fuentes et al., 2017). Regardless of the processes involved in the evolution of the margin during Triassic times (Mpodozis and Kay, 1990, 1992; del Rey et al., 2016; Coloma et al., 2017; Fuentes et al., 2017), the Las Tórtolas Formation and the other late Paleozoic accretionary complexes in north and central Chile constituted the continental basement of the arc magmatism related to the First Stage of the Andean cycle (Upper Triassic) (Charrier et al., 2007). 2.2. Local geology The Las Tórtolas Formation conforms a N-S-directed, around 12 kmwide strip of meta-volcanosedimentary rocks that corresponds to the late Paleozoic basement of the Upper Triassic-Lower Cretaceous Coastal Range batholith in northern Chile (26°-29°S) (Fig. 1b). The basement of the Las Tórtolas Formation is unknown, whereas Triassic volcanosedimentary deposits (Pan de Azúcar and Cifuncho formations) are unconformably deposited over it (Naranjo and Puig, 1984). The lithological and structural characteristics across the Las Tórtolas Formation

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linear and linear-planar fabric domains. A detailed structural description of the mélange unit in the Chañaral and Obispito areas (Fig. 1c) is presented in Section 3.1. In contrast to the other units defined as the Western Series of the late Paleozoic accretionary complexes of the Chilean coast (Fig. 1b), the mélange unit of the Las Tórtolas Formation does not show HP rocks, and the metamorphic temperature conditions are in the range of 350–400 °C, which coincides with the low-grade metamorphism proposed for the entire Las Tortolas Formation (Miller, 1970; Aguirre et al., 1972; Marioth, 2001; Navarro, 2013; Fuentes et al., 2016). Moreover, this domain has been related to the shallow zone of a subduction plate boundary or the basal zone of an accretionary complex according to its lithological and structural characteristics and tectonic correlations (Fuentes et al., 2016, 2017).

allow us to distinguish three domains, from west to east: the mélange unit, the volcanic domain, and the eastern metaturbiditic unit (Fig. 1c). It mostly consists of a monotonous sequence of interbedded or tectonically mixed sandstones and shales, with a few limestones, pelagic chert and conglomerates, and oceanic, mainly E-MORB, basaltic rocks metamorphosed to low-grade conditions ranging between greenschist facies (Miller, 1970; Aguirre et al., 1972) and pumpellyiteactinolite facies for the mélange rocks, and prehnite-actinolite facies for the eastern metaturbiditic unit (Marioth, 2001). The sedimentary structures and rock types indicate a deep-sea basin-plain depositional environment for the Las Tórtolas Formation (Bell, 1982). These synsedimentary structures are preserved in the eastern metaturbiditic unit but are obliterated by the intense deformation in the mélange unit. Late deformational processes and contact metamorphism are associated with the emplacement of Upper Triassic to Lower Jurassic intrusives and the activity of N-S and NW-SE directed faults associated with the Atacama Fault System during the First Stage of the Andean cycle (Fig. 1c).

2.2.2. Volcanic domain The volcanic domain of the Las Tórtolas Formation conforms an elongate unit showing N-S directed, E-dipping thrusted contacts with the western mélange unit and the eastern metaturbiditic unit (Figs. 1c, 2b). The volcanic domain consists of basaltic hyaloclastitic lavas to the bottom (Fig. 2c), overlain by thick lava flows interbedded with quartzites and phyllites, and pyroclastic brecciated deposits and rhyolites to the top. Except for the basal N-MORB hyaloclastites, most of the metabasites that comprise the volcanic domain yield E-MORB signatures and have been related to an enriched mantle source, such as a mantle plume. The basaltic blocks included in the mélange unit gave similar trace element contents than the E-MOR basalts analyzed in the volcanic domain (Fuentes et al., 2017). The structural characteristics and the low-grade metamorphism of the volcanic domain are similar to those observed in the eastern metaturbiditic unit of the Las Tórtolas Formation. The volcanic domain may be associated with an oceanic ridge or a group of seamounts

2.2.1. Mélange unit The mélange unit represents the westernmost area of the Las Tórtolas Formation. This domain is characterized by the intense disruption of the sedimentary sequence and is conformed by centimetric to decametric quartzite blocks in a phyllitic matrix (Fig. 2a). Less abundant blocks of limestone, pelagic chert, conglomerate, and metabasite can also be observed. Except for the metabasites, which are geochemically similar to those observed in the volcanic domain, the rest of the materials that comprise the mélange unit are identical to those found in the metasedimentary sequences of the eastern metaturbiditic unit. The distinctive structural feature of the mélange unit is the block-inmatrix texture resulted after an intense ductile deformation. The shape of the quartzite blocks and the phyllitic matrix textures, which show a constant N140°-150°E azimuth for the foliations and lineations, defines

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Fig. 2. (a) Chloritized metabasites as blocks in the phyllitic matrix of the mélange unit (b) The Infieles fault in the Chañaral area shows dextral reverse kinematics and put in contact the mélange unit (W) and the eastern metaturbiditic unit (E). (c) Hyaloclastites in the volcanic domain. (d) SW-vergent thrust-related folds are one of the most outstanding structural characteristics of the eastern metaturbiditic unit of the Las Tórtolas Formation.

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accreted to the frontal accretionary formations (eastern metaturbiditic unit) during early to middle Permian times (Fuentes et al., 2017). 2.2.3. Eastern metaturbiditic unit The eastern unit of the Las Tórtolas Formation shows regional correlations with the eastern low-grade metasedimentary domains of the Chilean late Paleozoic accretionary complexes (Fig. 1b), which are related to the frontal accretion in the convergent margin (e.g., Aguirre et al., 1972; Hervé, 1988; Willner, 2005; Willner et al., 2005, 2012). It consists of a monotonous sequence of interbedded sandstones (quartzites) and shales (phyllites), with a few limestones, pelagic cherts and conglomerates. In contrast to the mélange unit, the quartzite-phyllite sequence shows layers of great lateral continuity. As in the volcanic domain, the structure of the eastern metaturbiditic unit is characterized by N160°E to N-S trending, kilometric scale thrust faults and associated fault-propagation folds (Fig. 2d), which generated a penetrative, NE to E dipping, axial-plane foliation in the phyllite layers. Both thrusts and associated folds are W- to SW-vergent. A more detailed structural description is presented below. 3. Structural description 3.1. Mélange unit 3.1.1. Chañaral The structure of the tectonic mélange is defined by the shape of the essentially quartzitic blocks and the fabrics developed in the phyllitic matrix, which constitute linear (L) and linear-planar (L-S) domains in the Chañaral area. Along the L domains, the quartzite blocks show a constrictional shape while the foliation planes were not identified in the phyllite matrix, arranged as L-tectonites with “pencil structure”. In those domains, the long axes of quartzite ellipsoidal bodies and the phyllite lineations (L1) present NW-SE trends, plunging 5° to 20° to the NW and SE (Fig. 3). The less constrictional shape of the quartzite blocks characterizes the L-S domains, as well as the linear-planar fabric observed in the phyllitic matrix. The mechanism that acted in the mélange to yield the described structural configuration is boudinage, in addition to heterogenous shearing (S-C structures, Fig. 3). The tectonic foliation (S1) defined by the shape of the quartzite bodies and the phyllite cleavage shows a constant N130°-150°E azimuth, dipping 40° to 85° to NE and SW (Fig. 3). This variation in the dip direction of S1 reveals cartographic-scale folds that are close to tight, moderately inclined chevron folds according to the calculated statistical axial plane and the absence of gently dipping foliations. As few tight to isoclinal intrafolial folds (B1) were found in the Chañaral area (Fig. 3), and their fold hinges and axial planes coincide with the lineation (L1) and foliation (S1) defined in the mélange unit respectively, the large map-scale folds affecting S1 are considered as B2 folds. This second stage folding did not yield a second tectonic foliation, except rare N150°-160°E directed, NE-dipping crenulation cleavage planes (S2) associated with the axial planes of few, and very small kink folds. This is consistent with the fold morphologies expected for the high mélange viscosity contrast between the quartzite blocks and the phyllitic matrix (Ramsay and Huber, 1987). The variable plunging of L1 lineations might be partly caused by B2 folds, as will be analyzed later in this work. S-C, boudinage and pinch and swell structures are ubiquitous along the mélange unit and are considered here as the prime responsible for the intense disruption of the original sedimentary sequence (Fig. 3). These structures are characteristic of shear zones and, due to their asymmetric arrangement, can be used as kinematic criteria, attesting the activity of a non-coaxial flow component, and pointing to a topto-the-southwest kinematics for the deformation zone. The study of the shape preferred orientation (SPO) determined in the ellipsoidal quartzite blocks has shown that constrictional blocks predominate in the L domains, while those measured in the L-S domains plot in the transition between the constrictional and flattened shapes

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(Fig. 3). The high aspect ratios obtained in the mesoscopic fabric are not in agreement with the microscopic study (quartz microfabric), albeit the ellipsoid shapes are similar (Fig. 3). These scale-dependent differences in the observed SPO fabric intensity results may be related to the deformation mechanism, assigned to a process of dissolutionprecipitation creep that affected unconsolidated or poorly consolidated sediments during the mélange formation (Fuentes et al., 2016). 3.1.2. Obispito In the Obispito area (Figs. 1c, 4), the tectonic mélange shows structural characteristics, fabric domains, and foliation and lineation orientations similar to those in the northern Chañaral outcrops (compare Figs. 3 and 4). Again, L and L-S or S domains are determined by the constrictional or flattened shape of the quartzite blocks, and the Land S-tectonites developed in the phyllitic matrix (Fig. 4). Lineations (L1) show fairly constant trends (around N135°E and N320°E), with gentle to moderate plunges in the range between 5° and 35°. NW-SE trending foliations (S1) dip between 35° and 60° to the NE and SW describing almost upright map-scale B2 folds, similar to those observed in the Chañaral area (Fig. 4). However, some lithological and textural differences deserve to be noticed. To the SSW of the study area, the sandy beds are more abundant in the metaturbiditic sequence, which shows a fine (1 to 10 cm) and highly deformed interlayering between phyllitic and quartzitic beds. The intense stretching of the quartzite layers (mostly asymmetric pinch and swell structures and boudinage) triggered lateral thickness variations of the metasedimentary beds and they barely preserve their original continuity (Fig. 4). Besides, abundant tight folds are nucleated in specific areas of these sandy domains (B1 folds). Although they present different styles and dips of their axial planes, fold axes show the same NW-SE directed, gently plunging orientations than the L1 lineations of the mélange unit (Fig. 4). As in the Chañaral area, the mesoscopic fabric described by the SPO of the quartzite blocks reveals constrictional shapes for linear domains (L). However, the mélange unit in Obispito yields clearly flattened block shapes in the L-S and S domains (Fig. 4). Those results were mainly obtained in the more sand-rich, interlayered areas. The asymmetric structures in Obispito consistently reveal a top-to-thesouthwest kinematics for the entire area. 3.2. Volcanic domain and eastern metaturbiditic unit The structure of the Eastern Series and the volcanic domain is determined by the original depositional bedding (So) defined by the contacts between phyllite and quartzite layers of the metaturbiditic sequence, and among the distinct types of volcanosedimentary rocks, respectively. A highly penetrative tectonic foliation was mostly developed in the phyllite layers (S1). This axial plane foliation resulted from the W to SW vergent, inclined folds associated with the propagation of imbricate, N160°E directed thrust faults (Fig. 1c). This NNW-SSE orientation dominates both the volcanic domain and the eastern metaturbiditic unit, although small-scale variations are observed between NNW-directed and N-directed domains limited by major map-scale thrust faults, mostly along the eastern metaturbiditic unit. Besides, a central narrow strip of the metasedimentary formation shows upright to inclined E-vergent folds, and N-S trending, W-dipping thrust faults, and a small elongated N-S domain of mélange-like rocks that may be associated with a region dominated by retro-vergent structures (Fig. 1c) (e.g., Davis et al., 1983; Naylor et al., 2005). Boudinage, pinch and swell and necking structures were observed in a few localities in quartzite and volcanic layers. These structures were subsequently thrusted and folded during the intense ENE-WSW shortening that affected and determined the structural characteristics of the volcanic domain and the eastern metaturbiditic unit. This structural homogeneity between the oceanic rocks and the turbiditic sequences points to a common evolution of both units in the frontal area of the

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Fig. 3. Structural sketch, spherical projection and Flinn diagram showing the main structural characteristics of the mélange unit at the Chañaral area. Photographs show the essential structures and fabrics affecting the mélange unit. The spherical projections in this and the rest of figures have been made using the program Stereonet by Richard Allmendinger, based on the algorithms described in Allmendinger et al. (2012). Equal-area, lower-hemisphere projection. The measurements of the quartzite block axes used to obtain the Flinn shape parameters were carried out in the SPO stations located in the structural sketch. See Fuentes et al. (2016) to review the details of the quartzite microfabric.

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Fig. 4. Structural sketch, spherical projection and Flinn diagram showing the main structural characteristics of the mélange unit at the Obispito area. Photographs show the essential structures and fabrics affecting the mélange unit. S-tectonites, D1 and D2 fold images were taken in the fold nucleation areas. The measurements of the quartzite block axes used to obtain the Flinn shape parameters were carried out in the SPO stations located in the structural sketch.

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accretionary prism after the accretion of the mainly basaltic oceanic reliefs (Fuentes et al., 2017). 4. Application of the model of transpression with oblique extrusion Most of the structures described above for the Chañaral mélange (L1, S1, B1) have been ascribed to a first deformation phase, D1 (Fuentes et al., 2016). The kinematic analysis of those D1 structures is the main objective of this work, although the effects of latter deformation phases, particularly the D2 folding, are also taken into account. 4.1. Description of the model Kinematic models applied to deformation zones are useful to constrain the orientation and shape of the finite strain ellipsoid within these zones resulting in foliations, lineations and strain markers of tectonites (e.g., Ramsay and Graham, 1970; Sanderson and Marchini, 1984; Fossen and Tikoff, 1993). Scale-independent, extended models of transpression have surpassed their original specificity to become a standard tool to analyze 3D deformation zones of any size and orientation. Consequently, these 3D kinematic models have greatly contributed to the understanding of the dynamic processes acting in deformation zones in different tectonic settings (see review by Díaz-Azpiroz et al., 2018). Deformation zones are kinematically modeled by a velocity gradient tensor that mathematically describes the motion of particles within the zone along time (e.g., Fossen and Tikoff, 1993). Particle paths are theoretical lines that illustrate particle motion and the three flow apophyses (for 3D flows) separate fields with contrasting motions (Passchier, 1997). Integration with time of the velocity gradient tensor yields the evolution of finite strain geometry. Despite some exceptions (e.g., Robin and Cruden, 1994; Jiang, 2007), most 3D kinematic models assume steady state, homogeneous deformation affecting a tabularshaped zone with parallel boundaries. Additionally, permitting slip along fixed zone boundaries and a free surface solves the deformation compatibility problem. In this analysis, we use the model of triclinic transpression with inclined extrusion (Fernández and Díaz-Azpiroz, 2009) (Fig. 5b), which has been successfully applied to ductile (Fernández et al., 2013), essentially brittle (Díaz-Azpiroz et al., 2014) and even active (Alonso-Henar et al., 2018) deformation zones. The reference frame is fixed to the

(a)

deformation zone (X1 parallel to the strike of the zone boundaries, X2 normal to the boundaries, X3 parallel to the dip direction). The zone is initially considered vertical, such that for inclined natural deformation zones (like in the cases studied in this work), the model results are rotated accordingly. As in other models, the flow within the deformation zone is described by the simultaneous combination of two components, in this case, a non-coaxial simple shearing and a coaxial component. _ Its Simple shearing is characterized by the simple shear strain rate (γ). direction is contained within the deformation zone and may show any orientation (denoted by angle ϕ) respecting the zone strike. The coaxial component is defined by three mutually perpendicular stretching axes ( ε_ b is normal to the zone boundaries whereas ε_ a and ε_ c may show any orientation parallel to the zone boundaries, Fig. 5b). The stretching across the deformation zone (ε_ b ) can be either contractional or extensional, resulting in transpressional and transtensional zones, respectively. On the other hand, ε_ a and ε_ c can also be contractional or extensional, with ε_ a describing the maximum lengthening or the minimum contraction parallel to the zone boundaries. Different 3D kinematic situations may arise from the multiple combinations of these three coaxial components (extended transpression models of Fossen and Tikoff, 1998) described by the Flinn k value or the logarithmic K value (see e.g., Ramsay and Huber, 1983). The axis ε_ a forms an angle υ (extrusion obliquity for transpression) with the dip direction of the zone and a deflection angle ζ with the simple shear direction (Fig. 5b). The ratio of the simple shearing to the coaxial component of the flow is proportional to the sectional vorticity number (Wk, Truesdell, 1954), which is obtained from the vorticity tensor, a decomposed part of the velocity gradient tensor. Wk values range from 0 (only coaxial flow) to 1 (only simple shearing). Therefore, Wk describes the rotational compo! nent of the flow, which occurs around the vorticity vector (ω ), oriented parallel to the pole to the vorticity normal section (VNS, Robin and Cruden, 1994). The VNS theoretically coincides with the section showing maximum structural asymmetry in tectonites, and its orientation is a fundamental element of the deformation zone kinematics (DíazAzpiroz et al., 2018). The eigenvalues and the associated eigenvectors of the left Cauchy-Green deformation tensor obtained from the integration of the defined velocity gradient tensor are, respectively, the length (principal quadratic extensions λ1 ≥ λ2 ≥ λ3) and the orientations of the principal axes of the finite strain ellipsoid (X, Y, and Z, respectively) with respect to the external reference frame (Truesdell and Toupin, 1960). They can be compared with the strain fabric of the deformation zone,

X2

(b)

X2

c

X1 X3

X1

a

X3

a

b

Fd Fd Fig. 5. Schematic block diagrams illustrating the essential features of two selected models of transpression. (a) Classical monoclinic transpression model of Sanderson and Marchini (1984). (b) Model of triclinic transpression with oblique extrusion and a general coaxial component (Fernández and Díaz-Azpiroz, 2009). Explanation and description of variables in the main text.

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such that it is assumed that the lineation and the pole to the foliation are good proxies for the X and Z axes, respectively, in particular at high strains (e.g., Williams, 1976; Ghosh, 1982). From the latter, it turns out that the finite strain fabric of a deformation zone permits to constrain the corresponding flow parameters (ϕ, υ, ! ζ, Wk, ω , VNS, K). This results directly from the boundary conditions, such that the simple shear direction and Wk are a function of the angle ƒ! Θ between the pole to the deformation zone boundary and F d (Fig. 5), the displacement vector between converging blocks (e.g., Jiang and Williams, 1998; Schulmann et al., 2003; Díaz-Azpiroz ! et al., 2014). In turn, ω is parallel to the deformation zone boundaries and normal to the simple shear direction. Moreover, it has been sugƒ! gested that the orientation of F d is represented by the oblique flow apophysis of 3D monoclinic flows (Fossen and Tikoff, 1998). Therefore, relevant tectonic interpretations may arise directly from the outcome of 3D kinematic models. 4.2. Choice of the relevant kinematic parameters According to the geological characteristics described above, the Chañaral mélange can be considered as a unit resulting from deformation inside a subduction plate boundary. Therefore, the aim of this work is to constrain the kinematic parameters controlling the flow during D1 along such a type of deformation zone. Application of the general model of triclinic transpression with oblique extrusion to the case of the Chañaral tectonic mélange followed two stages. First, a careful analysis of the range of values considered for the governing flow parameters has been done, based on the geological characteristics of the studied zone. Finally, the protocol suggested by Fernández et al. (2013) to test the model against natural cases has been used. The average orientation of the deformation zone to be modeled is that of the contact between the mélange unit and the volcanic domain and eastern metaturbiditic unit. An average N-S strike has been selected for the Chañaral zone, albeit it locally varies from NW-SE to NE-SW (Fig. 1). On the other side, the contact is more NNW-SSE oriented in the central and southern part of the mapped area (Fig. 1), such that a

NW

259

N165°E strike has been determined for the Obispito zone. Dip of the contact is more difficult to constrain. Present-day attitudes show dips exceeding 40° E (cross section in Fig. 1). However, those dips are unrealistically large for the shallow part of a typical subduction plate boundary, where gentle dips are commonly reported (e.g., Shreve and Cloos, 1986; Kimura et al., 2012), and are likely due to later deformation phases (Fuentes et al., 2017). Therefore, more geologically realistic dips of 10° to 30° have been considered in this work. The orientation of the VNS has been directly obtained in the field through a comprehensive observation of the degree of structural asymmetry at all the available natural exposure surfaces (Fig. 6). The section of maximum asymmetry has a very constant orientation throughout the studied region, although it slightly differs from the Chañaral to the Obispito areas. In Chañaral the VNS is oriented N60°E/85°NW, roughly normal to the lineation (L1). Meanwhile, in Obispito the maximum asymmetry surfaces show a mean orientation of N75°E/90°. Accordingly, the pitch of the simple shear direction on the boundary of the deformation zone (i.e., the angle ϕ) varies from ~60° N (Chañaral) to 90° (Obispito) (Fig. 7). The relative displacement sense deduced from the orientation of the VNS and the observed asymmetry ranges from reverse-dextral to pure reverse movement. The orientation of the extrusion direction (angle υ) is more difficult to constrain. In order to simulate monoclinic transpressional flows, extrusion direction has been considered as either parallel or normal to the simple shear direction (angle ϕ) at each zone. A range of intermediate orientations for the extrusion directions has also been applied to reproduce triclinic transpression (Fig. 7). North-plunging extrusion directions have been excluded from the systematic analysis performed in this work because a preliminary checking of their results evidenced that they were unable to explain the structural features shown by the Chañaral mélange. Distinct logarithmic K values describing the relative values of the stretching axes of the coaxial component have been examined. Given the presence of strongly linear fabrics and constrictional SPOs in both studied zones, values of K ≤ 1 (plane strain or flattening histories) have been discarded. In fact, all the available analytical models of transpression systematically predict flattening finite strain ellipsoids when the coaxial component consists of pure shearing or

SE WSW

ENE

ENE

WSW

Fig. 6. Field photographs illustrating the relative symmetric (left: photograph and sketch based on it) and asymmetric (right column) structures observed at distinctly oriented surfaces of the Chañaral mélange. The sub-vertical surfaces oriented ENE-WSE show the maximum asymmetry and are considered as the best indicator of the vorticity normal section.

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Fig. 7. Equal-area, lower-hemisphere projections showing the orientation of the main variables of the used triclinic transpression model for the studied zones: (a) Chañaral zone; (b) Obispito zone. The numerical values are shown in the tables below the spherical plots. The orientation of the extrusion direction (υ) is shown according its definition in Fig. 5, and also as its pitch on the shear zone boundary (SZB). Description of the variables in the main text.

general flattening (e.g., Fossen and Tikoff, 1998; Lin et al., 1998; Fernández and Díaz-Azpiroz, 2009). Therefore, K values of 1.1, 2, and 5, considered as three realistic figures for a constraining coaxial component, have been tested here. Finally, and in absence of clear quantitative indicators of vorticity, a wide range Wk values has been applied (Fig. 7). 4.3. Results The results of the applied transpression model are very varied, given the large number of possible combinations of the values considered for the constraining parameters (Fig. 7). To select the optimal combinations it is necessary to compare the predictions of the model with the available natural data. The protocol presented by Fernández et al. (2013) considers three main steps. The starting point (step 1a) is the comparison of the trajectories of the principal axes of the theoretical finite strain ellipsoid (axes X and Z) with the orientation of the natural fabrics (L1 lineation and pole to S1 foliation, respectively). As the natural data have not allowed determination of finite strain ellipsoids (Flinn diagrams in Figs. 3 and 4 only represent SPO), application of step 1b - comparison between the orientation of the principal axes of theoretical and measured finite strain ellipsoids- has not been possible. For the same reason, use of step 2 of the cited protocol, comparing the ellipticity and orientation of the finite strain ellipse on the VNS with theoretical values, has proven to be unfeasible in this case. However, a semiquantitative implementation of step 3 yielded acceptable results: the measured SPO of quartzite blocks and quartz grains, plotted on a Flinn diagram, are compared with the theoretical shapes derived from the model for the finite strain ellipsoid. Although SPO fabric ellipsoid and finite strain ellipsoid can be distinct for the same rock volume, the good correspondence observed in the field between the shape of the SPO ellipsoids and the structural style of the distinct domains (L, L-S, S) suggests that application of step 3 is possible for the Chañaral mélange.

Steps 1a and 3 of the protocol have been applied to the data of the D1 deformation phase that affected the Chañaral mélange to constrain the values of the critical parameters of the triclinic model that best match the natural data. Only the results of steps 1a and 3 that yield compatible combinations of ϕ, Wk, υ, and K have been considered. As an example, Fig. 8 shows, for some selected cases, poor and good fits for both steps at the two studied zones. Only those cases with simultaneous good fit for steps 1a and 3 (best combined fit in Fig. 8) are considered as compatible solutions and allow constraining the values of the controlling parameters. It is necessary to mention that the evaluation of the fit for step 1a has mainly considered the average orientations of the measured L1 and S1 (see, e.g., the density diagrams of Fig. 8). The dispersion of those fabrics is later, due to the effects of the D2 folding stage. The reorientation of the early fabrics by D2 is analyzed at the end of this section. The results are similar for both zones (Fig. 9). The best fit is obtained for low dips (10°) of the deformation zone boundary, although larger values (up to 30°) are also possible in Obispito. The simple shear direction shows a large pitch to the N (60°-90°). The K value of the coaxial component is 5 for both studied zones, although lower values (K = 2) are also acceptable in Obispito. The pitch of the extrusion direction is consistently low to moderate (30°-40°) and directed to the S-SE. Unfortunately, it has not been possible to constrain the value of the vorticity number (Wk) in this case, although it must be lower than 0.7 in the Chañaral zone (coaxial-dominated flow in that zone). The results of the Chañaral zone have been taken as an example to compute the orientation of the flow apophyses and the passive line rotation patterns (Fig. 9b). As expected, flow apophysis 1 is the line attractor, and it coincides with the location of the extrusion direction (Fig. 9a, b). Surprisingly, the flow trajectories show that the oblique apophysis (flow apophysis 3 in Fig. 9b) is not the flow repulsor, but the flow transit. Instead, flow apophysis 2 (located at 90° from the attractor along the deformation zone boundary) acts as flow repulsor in this case.

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k

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2

2

k

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3

k

n=18

n=16

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k

k

1

1

2

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k

Fig. 8. Selected cases of comparison of the results of the model of triclinic transpression with the data of the Chañaral (a) and Obispito (b) zones. Left two columns: Equal-area, lowerhemisphere projection diagrams showing the comparison of L1 and poles to S1 with the theoretical X (small black geometric figures) and Z (small red geometric figures) axes of the finite strain ellipsoid for distinct vorticity numbers (step 1a). Density contours (shades of red for poles to S1, and shades of gray for L1) follow the method of Kamb (1959). Central column: Comparison of the best fit for steps 1a and 3 (explanation in the main text), with indication of the only compatible solution (red rectangle). Right columns: Comparison in the logarithmic Flinn diagram of the measured SPO fabrics with the predictions of the model of triclinic transpression for the shape of the finite strain ellipsoid (black continuous or discontinuous lines depending on the value of the vorticity number) (step 3).

The orientation of the tight to isoclinal B1 folds can theoretically be determined for the estimated parameters of the transpression model applied to the Chañaral mélange. It is first necessary to assume that the folded layers were initially subhorizontal, which seems very reasonable for the still undeformed sedimentary sequence. Then, the initial azimuth of the B1 fold hinges can be calculated as perpendicular to the trend of the maximum horizontal instantaneous shortening direction (e.g., Fossen et al., 2013). Afterwards, fold hinges are considered to rotate as passive lines (e.g., Fossen et al., 2013) and the locus occupied by the theoretical fold hinges is calculated and plotted against that of the measured B1 folds (Fig. 10a, b). Although the trend of the theoretical B1 fold hinges is identical to that of natural B1 folds, the latter show in general larger plunges. This feature is assumed here to derive from the effects of the D2 folding stage. Therefore, to analyze the reorientation of L1, S1, and B1 during D2, the methods of folding of inclined lines and planes explained by Ramsay (1967) have been followed. In all cases, the buckling mechanism of folding has been assumed for D2. Selected initial orientations of L1, S1 and B1 have been chosen from the predictions of the transpression model fitted to the natural data. According to the orientation of the B2 folds explained above (subhorizontal, NW-SE directed hinge lines, and steeply, NE-dipping axial surfaces) the theoretical reorientation of lines (L1 and B1) matches well the observed dispersion of lineations and fold hinges (Fig. 10c, d), although NW-plunging lines are not easily modeled for the Obispito zone. Also the theoretical rotation of the poles to S1 encompasses

most of the observed orientations (Fig. 10e, f), with some difficulties again to explain the steeply NE-plunging poles of the Obispito zone. Probably a more comprehensive treatment, with a wider selection of the initial orientations of the planar and linear fabrics would have allowed a complete explanation of the dispersion measured in the natural data. Also the effects of younger phases should be taken into account, as is acknowledged in the discussion section. Despite this, the explored cases offer a remarkable fit to the observed fabrics. In summary, we can reasonably conclude that the procedure followed for the adjustment of the Chañaral mélange to the triclinic transpression model is robust and offers reliable results for the tectonic interpretation of the unit. 5. Discussion 5.1. Kinematic implications of the results of the triclinic transpression model to the Chañaral tectonic mélange A relevant structural feature of the Chañaral mélange is the almost orthogonal arrangement between the VNS and L1. Indeed, this is not a surprising observation, given that modern kinematic models predict that, in the more general cases, the orientation of the VNS is independent from the finite strain fabric (see, e.g., a recent discussion of this topic in Díaz-Azpiroz et al., 2018). The natural case studied in this work reinforces the idea that the VNS is an essential element of

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Fig. 9. Equal-area, lower-hemisphere projection diagrams summarizing the results of application of the model of triclinic transpression to the structural data of the Chañaral mélange. (a) Orientation of the main parameters controlling the model (see the main text for further explanation and discussion). The table below the spherical projections shows the numerical values or range of values for each parameter resulting from the application of the model at the Chañaral and Obispito zones. (b) Example of passive line rotation pattern for the Chañaral zone (υ = 130°, Wk = 0.35). Note how trajectories depart from flow apophysis 2 (repulsor) and converge towards flow apophysis 1 (attractor, extrusion direction).

deformation zones, and that it must be determined independently, without previous assumptions about its hypothetical relationship with planar and linear fabric elements associated with the finite strain. The complex triclinic transpression model described in this work, and satisfactorily applied to the case of the Chañaral mélange, does not follow the commonly assumed condition that the convergence diƒ! rection ( F d and angle Θ, Fig. 5) between the blocks separated by the deformation zone coincides with the oblique flow apophysis (Fossen and Tikoff, 1998). The convergence direction should match instead the trend of the flow apophysis 2 that has been determined here as N50°60°E (Fig. 9a). Considering the interpretation of the Chañaral mélange as the consequence of deformation within a subduction plate boundary or subduction channel (as described for the Franciscan subduction complex in a recent review by Krohe, 2017), this convergence direction can be taken as a proxy to the relative velocity vector between the involved oceanic and continental plates during the late Paleozoic, at least at the vicinity of this plate boundary. This estimated convergence direction agrees with previous results determined in different parts of the Eastern and Western Series based on independent structural, sedimentary and tectonic criteria (e.g., Bell, 1984, 1987; Bahlburg, 1987; Willner, 2005; Kleiman and Japas, 2009; García-Sansegundo et al., 2014; Hyppolito et al., 2014, 2015; Fuentes et al., 2016, 2017; Creixell et al., 2016). On the other side, the estimated extrusion direction shows a low pitch on the boundary of the deformation zone (Figs. 9, 11). This feature implies that the flow kinematics imposed particle paths that led to low to moderate downward displacements of material, while its transportation sub-parallel to the direction of the plate margin was considerably larger (Fig. 11). Therefore, the kinematics of the deformation zone (here interpreted as a subduction channel) that generated the Chañaral mélange prevented upward displacement of deep HP rocks of the accretion complex. This agrees with the metamorphic conditions recorded at Chañaral (greenschist to pumpellyite-actinolite facies without evidences of HP metamorphism: Miller, 1970; Aguirre et al., 1972;

Marioth, 2001; Navarro, 2013; Fuentes et al., 2016), and strongly contrasts with the presence of HP relics in other segments of the late Paleozoic Chilean margin (e.g., Bahlburg et al., 2009; Hervé et al., 2013; Pankhurst et al., 2016). This topic will be addressed in more detail later in this discussion. Our results illustrate a complex strain-partitioning pattern produced at different scales. At the smaller scale, the comparison between the measured SPO and the theoretical ellipsoids computed by the model (Figs. 3, 4, and 8) permits to suggest that Wk varied spatially, such that the L domains correspond with the bands with lower Wk values (b0.7), while the L-S or S domains appear where Wk values exceeded 0.7 (Fig. 11). Given the constant value across both studied zones of the rest of the kinematical controlling parameters (ϕ, υ, K), the spatial variation of vorticity is indicating a type of deformation partitioning, similar to what it is observed at other shear zones or at meso- and microscopic scales (e.g., porphyroclast systems, S-C or S-C′ composite planar fabrics, etc.). In complex 3D shear zones, this vorticity partitioning is typically related to local rheological differences and/or to the heterogeneous geometry of the zone (e.g., Sullivan and Law, 2007; Sullivan, 2008; Barcos et al., 2015). Deformation partitioning occurs also along the subduction channel, as evidenced from different ϕ values in Chañaral (ϕ = 60°) and Obispito domains (ϕ = 90°), likely due to slight variations in the orientation of the deformation zone (N0°E/10°E in Chañaral, N165°E/ 10–30°E in Obispito) and a common convergence vector. Similar along-strike deformation partitioning has been documented in other shear zones (e.g., Barcos et al., 2015). At an even larger scale, the deformation partitioning is also observed when comparing the structural characteristics of the Chañaral mélange with those of the volcanic domain and eastern metaturbiditic unit (Figs. 1, 3, 4). The N-S- to NNWSSE-oriented structures (folds, thrusts, and fabrics) in the latter two domains contrast with the NW-SE fabrics measured in the mélange. The azimuth shown by the boundaries between the L and L-S or S domains in the mélange, however, is sub-parallel to the structures in the eastern

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(b)

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(c)

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L1 L1 D2

(e)

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N

S1

S1 Poles to S1

Poles to S1

E=3 Contour interval=2 Fig. 10. Equal-area, lower-hemisphere projection diagrams showing the simulation of the D1 and D2 folding events according to the results of the applied model of triclinic transpression. (a) and (b) Theoretical orientation of the B1 fold hinges (blue-shaded areas) computed from the results of the transpression model for the Chañaral and Obispito zones, respectively. Slightly distinct orientations result depending on the selected value of υ. Also represented are the natural B1 hinges measured at each zone (orange squares). (c) to (f) Variation of the orientation of the X (c and d) and Z (e and f) axes of the finite strain ellipsoid as a consequence of the D2 folding event, and for selected predictions of the applied transpression model. Left: Chañaral zone; right: Obispito zone. Triangles: initial position of the line; circles: progressive reorientation of the line in the limbs of the D2 folds; the yellow area represents the range of possible solutions between those predicted for the selected cases). Contour densities for L1 (gray shades) and poles to S1 (red shades) follow the method of Kamb (1959) (see Fig. 9).

regions. Finally, the structures mapped in the volcanic and eastern metaturbiditic unit are kinematically consistent with the simple shear direction determined in the mélange unit, whose azimuth ranges from N60°E to N80°E (e.g., Fig. 9). It can be suggested that the simple shear component associated with the plate convergence was particularly localized at the accretionary complex, with emplacement directions forming a high angle to the azimuth of the plate boundary. Similar kinematics can also be observed at the mélange along the domains with higher vorticity number, although the supposed subduction channel

seems to have focused most of the coaxial component due to the plate oblique convergence. Interestingly, the D2 folding phase in the Chañaral mélange is almost coaxial with the older structures (L1 and B1), producing on them only minor reorientations mainly affecting the plunge sense or value (Fig. 10). This seems to indicate that the essentials of the kinematic pattern did not vary from the first to the second deformation phase. It can be suggested that the blocking of the supposed subduction channel, either due to rheological or tectonic causes as discussed below in this

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a

c a

b

k

k

Fig. 11. Schematic block diagram depicting the main results of the application of the model of triclinic transpression to the Chañaral mélange. The inset explains the meaning of the main parameters controlling the kinematic evolution of the zone, and it coincides with Fig. 5b conveniently rotated to coincide with the orientation of the supposed subduction channel.

section, led to the deactivation of the structures that had previously absorbed the deformation within the channel. The persistence of the convergence direction allowed the development of folds coaxial with the previous structures. Nevertheless, part of the dispersion in the L1, S1 and B1 data observed in the spherical plots (Figs. 3, 4, 8, 10) might be ascribed to a subsequent N60-70E trending folding event that affected this area and may be related to the Jurassic NW-SE shortening characteristic of the First Stage of the Andean cycle in the continental block of the convergent margin (Scheuber and González, 1999; Ring et al., 2012). The applied kinematic model successfully explains many of the main structural characteristics of the Chañaral mélange. However, the limitations of the model should also be taken into account. As indicated above, the model assumes steady-state flow affecting tabular-shaped zones, slipping boundaries of the deformation zone, and homogeneous deformation. The unsteady flow is probably the rule in Nature (e.g., Jiang, 1994), although the mathematical complexities and the inherent variability of this particularly complex type of flow greatly hinder its application to natural examples. Moreover, as suggested by the described transition from D1 to D2, it is likely that the orientation of both the deformation zone at the plate boundary (supposed subduction channel) and the convergence velocity vector remained constant during the analyzed period. Therefore, it is reasonable to assume steady kinematic boundary conditions in this case. Slipping boundaries do not seem to be a major problem for the zone studied in this work. The upper boundary of the Chañaral mélange is marked by a major fracture, the Infieles

fault (Fuentes et al., 2016). Its lower boundary is not exposed, although a rather progressive gradient towards the oceanic crust is expected (Fig. 11). The model of triclinic transpression has been successfully applied to similar asymmetric shear zones with only one slipping boundary (e.g., the Southern Iberian shear zone, Fernández et al., 2013). The discretization of a heterogeneously deformed zone into small regions where homogeneous deformation can be reasonably assumed is a common practice in Structural Geology (e.g., Fossen and Tikoff, 1998). Actually, the explanation given here to the L, L-S and S domains observed in the studied regions attests how the model can satisfactorily deal with the problem of deformation heterogeneity. Other limitations arise from the particular application of the model to the studied zone. The impossibility of measuring ellipsoids of the finite deformation, greatly a consequence of the absence of reliable markers and of the deformation mechanisms operating in the mélange, reduces or eliminates the ability to apply several of the protocol steps to the studied rocks. Consequently, the values of some of the critical kinematic parameters of the model, in particular the vorticity number, have not been adequately constrained. Fortunately, the SPO fabrics offered a rough, semi-quantitative approximation to the shape of the finite strain ellipsoid, thereby allowing reasonable comparisons with the predictions of the model. Deformation phases younger than D1 (e.g., D2) have modified the orientation of the original structures. Part of the resulting dispersion in the orientation of L1, S1, and B1 has been adequately analyzed in this work with conventional structural techniques. The dip of the boundaries of the deformation zone was also probably modified after D1, but the applied model

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has considered a geologically reasonable range of dip angles. However, it cannot be ruled out that future detailed studies of those late phases may lead to minor modifications in the results of this work.

5.2. Applicability to the study of mélanges in subduction channels It seems relevant to ask here, using the Occam's razor, why it is necessary to apply such a complex model to any particular deformed zone. The answer is that things must be considered as simple as possible, but no simpler. Monoclinic models of transpression cannot explain the intricacies of the deformation in the Chañaral mélange, including the obliquity between the extrusion and simple shear directions, or the highangle that the VNS forms with L1. Also any transpression model with pure shearing as coaxial component fails to simulate the generation of constrictive fabrics (e.g., Sullivan and Law, 2007; Fernández and DíazAzpiroz, 2009). However, it is more instructive now to compare the implications of such a model with the simpler, 2D, classical kinematic models of tectonic mélanges generated in subduction channels. It is commonplace to find sketches of accretionary prisms and subduction channels emphasizing the structure, textures, metamorphic conditions or particle paths in 2D sections normal to the azimuth of the plate boundary (e.g., from the pioneering works of Cloos, 1982; Cloos and Shreve, 1988a, 1988b; to the more recent contributions of Kimura et al., 2012). To be aware of the implications of that standard procedure, it is illustrative to clarify that in those 2D cases the convergence direction ƒ! (vector F d ) must be normal to the azimuth of the plate boundary (Θ = 90°), which leads to pure dip-slip displacement for the simple shear component (ϕ = 90°), while the extrusion direction (ε_ a) is parallel (υ = 0°) to the dip-direction of the boundaries of the subduction channel (Fig. 12a). As a consequence, the kinematics of flow tends to be monoclinic and flow trajectories remain in the plane of the section. However, Philippon and Corti (2016) have shown that only 8% of

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plate boundaries show sub-orthogonal (Θ ≈ 90°) plate displacements. Oblique tectonics (Θ ≠ 90°) is the rule at convergent margins and, as a result, it is expected obliquity of the simple shear component (0≤| ϕ|≤90°), and oblique extrusion (υ ≠ 0°). Therefore, under oblique tectonics, particle paths crosscut the cross section normal to the azimuth of the plate boundary (Fig. 12b), and a 2D approximation is no longer acceptable. This is indeed the case for the zones of the Chañaral mélange studied in this work. Taking as a reference the results of application of the transpression model to the Chañaral and Obispito zones (Fig. 9), and assuming they are due to deformation inside a subduction channel, it is possible to estimate the amount of displacement produced by the coaxial component (following a procedure similar to that explained by Schulmann et al., 2003), and assuming a reference level of zero displacement coincident with the “control point” (a minimum in local capacity of the subduction channel, according to Cloos and Shreve, 1988a). Distance of the control point from the inlet (plate boundary) commonly varies from around 30 to 300 km (Shreve and Cloos, 1986). It is not possible to determine the position of the Late Paleozoic control point for the Chañaral mélange, and a distance of 30 km has been selected. Results for larger distances are simply the corresponding multiples of the values obtained for 30 km. Another necessary and unknown parameter is the strain rate. According to geological, geophysical and tectonic criteria (e.g., Pfiffner and Ramsay, 1982; Campbell-Stone, 2002; Zhu et al., 2006) a range of longitudinal strain rates normal to the boundaries of the subduction channel of 10−15–10−16 s−1 is considered as reasonable. The uncertainty in the duration of the accretionary process makes advisable to present the results for an arbitrary period as reference (in this case, 20 Myr). The considered strain rates yield particle displacements along the extrusion direction (i.e., upwards and NW-directed, oblique to the margin) ranging from up to 400 km (10−15 s−1) to b20 km (10−16 s−1) after 20 Myr (Fig. 12c, left diagram). However, the displacements along the dip of the deformation zone boundary are negative (downwards) and small or moderate, attaining b20 km in 20 Myr and

Fig. 12. (a) and (b) Schematic block diagrams showing the main kinematic characteristics of a subduction channel subjected to orthogonal (a) and oblique (b) convergence. (c) Representation of the main implications of the results of the applied triclinic transpression model in what concerns to the particle displacements associated with the coaxial component of flow. The left/right diagrams represent the displacements along the extrusion direction, and the dip of the deformation zone, respectively. (d) Representation of the displacement associated with the simple shear component. See the main text for further explanations.

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for the faster strain rate. This last result explains the lack of evidences of HP metamorphism in the Chañaral mélange, as mentioned earlier in this work. With the necessary caution due to the uncertainty in the aforementioned parameters, the results presented in Fig. 12c indicate an important tectonic flux of matter in a direction subparallel to the plate boundary, which is something rarely described previously (although lateral extrusion under transpression has been illustrated by e.g., Jones et al., 1997). The Chañaral mélange is a good example of how 3D deformation controls some of the main processes involved in complex tectonic settings (in this case, metamorphism in subduction). Thus, the profound knowledge of the kinematics provided by analytical models is of great help to better understand any tectonic setting. Analogously, the displacement associated with the simple shear component can be approximately estimated from the predictions of the used transpression model. The maximum values are obtained for Wk = 1, and range from around 5–40 km in 20 Myr (Fig. 12d), depending on the shear strain rate. More reasonable vorticity numbers (≤0.9) yield values lower than 15 km. The corresponding velocities for the simple shear component are, therefore, much b2 mm/yr. This figure should not be considered as an approximation to the relative plate velocity, because it does not take into account the coaxial component, nor the effects of the deformation partitioning inside and outside the subduction channel. This work highlights the need for a complete treatment of the kinematics of any given tectonic mélange. A simple 2D approach is probably not enough in most cases to satisfactorily explain the structural and metamorphic characteristics of the studied mélange. Alternatively, application of the methodology outlined in this work to distinct points along the same accretionary complex can serve to understand their tectono-metamorphic differences, and to identify and better describe the longitudinal variation along old plate boundaries of first-order tectonic parameters (like, for instance, the convergence direction). In any case, it must be remarked that kinematic modeling can only be applied to well delimited cortical domains where the observed structures can be interpreted in terms of finite strain resulting from a single bulk strain, and thus caution is needed to interpret the obtained results. Strain partitioning between adjacent domains at the same scale showing distinct kinematics, such as that observed between L and L-S/S domains in the Chañaral mélange, is common under different deformation conditions (e.g., Sullivan and Law, 2007; Fernández et al., 2013; Barcos et al., 2015). Strain partitioning typically occurs also at different scales, such that a given bulk strain is partitioned into different smaller domains, each accommodating a different component of the imposed flow (Lister and Williams, 1983; Tikoff and Teyssier, 1994; Schulmann et al., 2003; Díaz-Azpiroz et al., 2014). Therefore, the kinematics obtained from an analyzed domain cannot be directly extrapolated to similar domains in the area or to larger domains to account for tectonic interpretations (Carreras et al., 2013). Instead, detailed and multi-scale kinematic analyses supported by kinematic models permit to identify different strain partitioning patterns (e.g., Czeck and Hudleston, 2003; Sullivan, 2008; Fernández et al., 2013; Barcos et al., 2015; Nabavi et al., 2017) and provide relevant information applicable to large-scale tectonic interpretations.

5.3. Geotectonic implications in the Late Paleozoic accretionary process in the Chilean margin 5.3.1. The Chañaral mélange in the context of the late Paleozoic accretionary complexes of the Chilean coast The textural and structural characteristics of the mélange unit studied in Chañaral and Obispito outcrops are in agreement with those described for the tectonic mélanges. Therefore, the intense disruption of the original turbiditic sequence is determined by the activity of a reverse and dextral deformation zone. Material flow trajectories allowed those deformed sediments to be underplated at the base of the accretionary

wedge (Figs. 9, 11, 12), a plausible mechanism for the origin of the Chañaral mélange previously proposed by Bell (1987). The Las Tórtolas Formation has been traditionally related to a Carboniferous accretionary complex active in the southwestern margin of Gondwana after the accretion of several terranes during Paleozoic times (Bell, 1982, 1987; Charrier et al., 2007; Fuentes et al., 2016; Heredia et al., 2018). Accordingly, it is correlated with other coeval accretionary complexes that crop out along the Chilean coast (Fig. 1b). All these accretionary complex fragments are characterized by a western highly deformed domain (Western Series) and a metaturbiditic low-grade eastern domain (Eastern Series). Detrital zircons obtained in the Las Tórtolas Formation (included nonpublished data) yield maximum depositional ages of 360–350 Ma, which coincides with the ages obtained for the coastal accretionary complexes (26°-39°S) (Pankhurst et al., 2016; Hervé et al., 2013). Detailed structural analyses, like those presented in this work, are rare in those formations (e.g. Richter et al., 2007), where the structural descriptions are focused on revealing the deformational phases. However, main structural features as the W to SW vergence of the D1 phase are similar along the Western and Eastern Series. Other textural and lithological characteristics, as the presence of E-MORB and minor N-MORB oceanic rocks in the mélange unit and the volcanic domain reinforce this N-S correlation. Therefore, the main difference between the mélange unit of the Las Tórtolas Formation and the other late Paleozoic basal accretionary units is the absence of high-pressure rocks that are part of the blockin-matrix units. According to the results of the kinematic model tested in the Chañaral mélange (in Chañaral and Obispito localities), the oblique subduction generated flow trajectories at high angles relative to the dip direction of the contact plane between the convergent plates. This indicates that the material particles introduced in the deformation zone at the plate boundary (supposed subduction channel) covered longer distances along the strike of the plane than in its dip direction, before to became underplated to the basal accretionary complex, which implies that those rocks could not have been subjected to highpressure conditions. Similar differences were observed in the Franciscan subduction complex (USA) (Cloos, 1982). These distinctive characteristics of the Chañaral tectonic mélange are a consequence of the oblique convergence and might be explained by the local trend of the plate margin and the location of the northern Chilean margin in relation to the plate rotation pole (Euler pole). Indeed, a change in the margin azimuth from N-S to NW-SE has been proposed in the southwestern margin of Gondwana south of 34°S (Martin et al., 1999; Kleiman and Japas, 2009). The configuration of the oceanic plate or the presence of a spreading center that generated different convergence directions may be also considered. The notable differences observed in the final arrangement of the Western and the Eastern Series and oceanic basalts in the Chilean accretionary complexes may be also associated with the second deformation phase, related to the arrival of oceanic reliefs or seamounts to the margin (the volcanic domain in the Las Tórtolas Formation) (Hyppolito et al., 2015; Heredia et al., 2018; Díaz-Alvarado et al., 2019). 5.3.2. Deformational stages and the final configuration of the Las Tórtolas Formation The succession of deformational processes that finally conformed the lithological and structural zonation of the Las Tórtolas Formation between the Carboniferous and the middle Permian was determined by the NE- to ENE-directed subduction and the accretion in the western margin of Gondwana. According to the kinematic characteristics of the deformation zone established in the subduction plane boundary (supposed subduction channel) and the basal area of the accretionary prism, the block-in-matrix textures arose from the original turbiditic sequence producing a tectonic foliation (S1) and lineation (L1) in the phyllitic matrix and flattened and constrictional shaped blocks depending on the textural domain where they were formed. The limits between L and L-S domains are almost parallel to the plate boundary and were

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promoted by the spatially uneven predominance of the simple shear or coaxial component within the deformation zone. The main mechanism that acted in the mélange unit to yield the described structural configuration is boudinage, in addition to heterogenous shearing (SW-vergent, S-C structures) and folding processes that generated intra-foliation tight to isoclinal folds (Figs. 3 and 4). Those structures have been ascribed to a first deformation phase, D1. The foliation describes map-scale chevron folds (D2) after the overturn of the originally NE-dipping S1. The variable plunging of L1 lineations might be also caused by D2 folds (Fig. 10). This is consistent with the folding morphologies expected for the high mélange viscosity contrast between the quartzite blocks and the phyllitic matrix (Ramsay and Huber, 1987). According to the field measurements and the statistical results, the D2 folds in the mélange unit are homoaxial with the structures formed during D1. The thrust and associated fold structure that characterize the volcanic domain and the eastern metaturbiditic unit means that both domains were deformed together after the accretion of the oceanic rocks to the margin. As the former subduction and the accretion shared the convergence vector, previous structures related to the frontal area of the accretionary prism were reactivated during the arrival of the oceanic reliefs to the margin without any interference of structures. The identical geochemical and isotopic signatures observed in the basaltic rocks of the volcanic domain and the metabasite blocks of the mélange unit suggest a common enriched mantle source for the metabasites accreted to the margin from its base and from its front (Fuentes et al., 2017). The accretion of these oceanic rocks was dated between 300 and 270 Ma in the southern fragments of the accretionary complex (Hyppolito et al., 2014, 2015; García-Sansegundo et al., 2014), when the accretionary processes are expected to conclude along the margin. Therefore, the accretion of the volcanic domain and related rocks to the frontal area of the prism may have blocked the deformation zone at the subduction plate boundary, promoted the generation of the D2 folds in the mélange unit, and finally conformed the observed W-E zonation of the Las Tórtolas Formation. 6. Conclusions The predictions of the model of triclinic transpression with oblique extrusion (Fernández and Díaz-Azpiroz, 2009) have been checked against the structural features of the first deformation phase (D1) that affected the Chañaral tectonic mélange (Coastal Cordillera, northern Chile). That tectonic mélange is here interpreted as generated at a subduction channel active during the late Paleozoic in the southwestern margin of Gondwana. The boundaries of the supposed subduction channel are oriented N-S to NNW-SSE, and dip to the E. Two study areas (Chañaral and Obispito) were selected for this work, although scarce differences have been observed in both regions. A reasonable fit between model predictions and natural data has been obtained that allows constraining the range of possible values of the parameters controlling the flow kinematics. The vorticity normal section (section of maximum structural asymmetry) is observed at planes oblique or forming high angles to L1. The simple shear direction formed high pitch angles (60°N to 90°) in the deformation plane, evidencing reverse-dextral to pure reverse movement senses. The conspicuous constrictional fabrics observed in the mélange have been explained with constrictional coaxial components (logarithmic K values of 2 to 5). The extrusion direction shows a low pitch (30°-40°) to the south. The convergence direction between the blocks separated by the deformation zone is located at N50°-60°E, i.e., oblique to the azimuth of the plate boundary. Seemingly, the vorticity varied spatially, which accounts for the spatial distribution of L (lower relative vorticity numbers), and L-S or S (higher relative vorticity numbers) domains within the mélange. The model yields low to moderate pitchs of particle displacements downdip the deformation zone, in agreement with the low-pressure metamorphic conditions recorded in the mélange. The

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long relative particle displacements parallel to the direction of the deformation zone highlight the needs for a 3D study of deformation zones at subduction plate boundaries and subduction channels. The obtained oblique convergence is probably a consequence of the location of the northern Chilean margin relative to the plate rotation pole or other factors arising from the large-scale characteristics of the involved converging plates. A second phase of folding (D2) affected the structures and fabrics due to D1. It is here suggested that the deformation zone at the plate boundary became blocked during D2 by the arrival of submarine reliefs, without a concomitant change in the convergence direction between the involved plates and lithospheric blocks. The results of this work, with all the limitations inherent in the comparison between kinematic models and nature, highlight the importance of 3D studies in tectonic mélanges. A sound comparison between adjacent zones is also expected from this type of structural approach, with important implications in the tectonic understanding of old, fossil orogens.

Acknowledgements This work is part of the PhD study of Paulina Fuentes, carried out at the Department of Geology of the University of Atacama (Chile) and the University of Freiberg (Germany). The study has been funded with FONDECYT Project No 11140722 of CONICYT, and with the fund support of DIUDA 2013-22268 and DIUDA 2014-22273 projects (University of Atacama). The financial support of projects CGL-2013-48408-C3-1-9 and CGL2013-46368-P by the Spanish Ministry of Economy and Competitiveness is gratefully acknowledged. We gratefully acknowledge the editorial assistance and comments of the guest editor Andrea Festa, and the detailed reviews by two anonymous reviewers that greatly aided to improve the quality of this work. References Abbate, E., Bortolotti, V., Passerini, P., 1970. Olistostromes and olistoliths. Sedimentary Geology 4, 521–557. Aguirre, L., Hervé, F., Godoy, E., 1972. Distribution of metamorphic facies in Chilean outline. Krystallinikum. 9, pp. 7–19. Allmendinger, R.W., Cardozo, N.C., Fisher, D., 2012. Structural Geology Algorithms: Vectors & Tensors. Cambridge University Press, Cambridge, England (289 pp.). Alonso-Henar, J., Fernández, C., Martínez-Díaz, J.J., 2018. Applying the model of triclinic transpression with oblique extrusion to an active fault: the Alhama de Murcia Fault. Preliminary results. Resúmenes de la 3ª Reunión Ibérica sobre Fallas Activas y Paleosismología, Alicante, España (in press). Bahlburg, H., 1987. Sedimentology, petrology and geotectonic significance of the Paleozoic flysch in the Coastal Cordillera of northern Chile. Neues Jahrbuch für Geologie und Paläontologie, Monatshefte. 9, pp. 527–559. Bahlburg, H., Hervé, F., 1997. Geodynamic evolution and tectonostratigraphic terranes of NW-Argentina and N-Chile. Geological Society of America Bulletin 109, 869–884. Bahlburg, H., Vervoort, J.D., Du Frane, S.A., Bock, B., Augustsson, C., Reimann, C., 2009. Timing of crust formation and recycling in accretionary orogens: insights learned from the western margin of South America. Earth-Science Reviews 97 (1), 215–241. Barcos, L., Balanyá, J.C., Díaz-Azpiroz, M., Expósito, I., Jiménez-Bonilla, A., 2015. Kinematics of the Torcal Shear Zone: Transpressional tectonics in a salient-recess transition at the northern Gibraltar Arc. Tectonophysics 663, 62–77. https://doi.org/10.1016/j. tecto.2015.05.002. Bell, C.M., 1982. The lower paleozoic metasedimentary basement of the coastal ranges of Chile between 25°30′ and 27°S. Revista Geologica de Chile 17, 21–29. Bell, C.M., 1984. Deformation produced by the subduction of a Paleozoic turbidite sequence in northern Chile. Journal of the Geological Society, London 141, 339–347. Bell, C.M., 1987. The origin of the upper Palaeozoic Chanaral mélange of N Chile. Journal of the Geological Society, London 144, 599–610. Bettelli, G., Vannucchi, P., 2003. Structural style of the offscraped Ligurian oceanic sequences of the Northern Apennines: new hypothesis concerning the development of mélange block-in-matrix fabric. Journal of Structural Geology 25, 371–388. Braun, J., 1993. Three-dimensional numerical modeling of compressional orogenies: thrust geometry and oblique convergence. Geology 21, 153–156. https://doi.org/ 10.1130/0091-7613(1993)021b0153:TDNMOCN2.3.CO;2. Camerlenghi, A., Pini, G.A., 2009. Mud volcanoes, olistostromes and Argille scagliose in the Mediterranean region. Sedimentology 56, 319–365. Campbell-Stone, E., 2002. Realistic geologic strain rates. Geological Society of America, Annual Meeting 2002, Denver, CO https://gsa.confex.com/gsa/2002AM/ finalprogram/abstract_39029.htm. Carreras, J., Cosgrove, J.W., Druguet, E., 2013. Strain partitioning in banded and/or anisotropic rocks: implications for inferring tectonic regimes. Journal of Structural Geology 50, 7–21.

268

P. Fuentes et al. / Gondwana Research 74 (2019) 251–270

Casas, A.M., Gapais, D., Nalpas, T., Besnard, K., Román-Berdiel, T., 2001. Analogue models of transpressive systems. Journal of Structural Geology 23, 733–743. Charrier, R., Pinto, L., Rodríguez, M.P., 2007. Tectonostratigraphic evolution of the andean orogen in Chile. In: Moreno, T., Gibbons, W. (Eds.), The Geology of Chile. Geological Society, London, pp. 21–114. Czeck, D.M., Hudleston, P.J., 2003. Testing models for obliquely plunging lineations in transpression: a natural example and theoretical discussion. Journal of Structural Geology 25, 959–982. Cloos, M., 1982. Flow mélanges: Numerical modeling and geologic constraints on their origin in the Franciscan subduction complex, California. Geological Society of America Bulletin 93, 330–345. Cloos, M., Shreve, R.L., 1988a. Subduction-channel model of prism accretion, mélange formation, sediment subduction, and subduction erosion at convergent plate margins: 1. Background and description. Pure and Applied Geophysics 128, 455–500. Cloos, M., Shreve, R.L., 1988b. Subduction-channel model of prism accretion, mélange formation, sediment subduction, and subduction erosion at convergent plate margins: 2. Implications and discussion. Pure and Applied Geophysics 128, 501–545. Coloma, F., Valin, X., Oliveros, V., Vásquez, P., Creixell, C., Salazar, E., Ducea, M.N., 2017. Geochemistry of Permian to Triassic igneous rocks from northern Chile (28°-30°15′ S): Implications on the dynamics of the proto-Andean margin. Andean Geology 44 (2), 147–178. Cousineau, P.A., 1998. Large-scale liquefaction and fluidization in the cap chat mélange, Quebec appalachians. Canadian Journal of Earth Sciences 35 (12), 1408–1422. Cowan, D.S., 1985. Structural styles in Mesozoic and Cenozoic mélanges in the western Cordillera of North America. Geological Society of America Bulletin 96, 451–462. Cowan, D.S., Brandon, M.T., 1994. A symmetry-based method for kinematics analysis of large slip brittle fault zones. American Journal of Science 294, 257–306. Creixell, C., Oliveros, V., Vásquez, P., Navarro, J., Vallejos, D., Valin, X., Godoy, E., Ducea, M.N., 2016. Geodynamics of Late Carboniferous Early Permian forearc in North Chile (28°30′-29°30′ S). Journal of the Geological Society 173 (5), 757–772. Davis, D., Suppe, J., Dahlen, F.A., 1983. Mechanics of fold-and-thrust belts and accretionary wedges. Journal of Geophysical Research 88 (B2), 1153–1172. Díaz-Alvarado, J., Galaz, G., Oliveros, V., Creixell, C., Calderón, M., 2019. Fragments of the late Paleozoic accretionary complex in central and northern Chile: similarities and differences as a key to decipher the complexity of the pre-Andean cycle. In: Folguera, A., Horton, B. (Eds.), Andean Tectonics. Elsevier (in edition). Díaz-Azpiroz, M., Barcos, L., Balanyá, J.C., Fernández, C., Expósito, I., Czeck, D.M., 2014. Applying a general triclinic transpression model to highly partitioned brittle-ductile shear zones: a case study from the Torcal de Antequera massif, external Betics, southern Spain. Journal of Structural Geology 68, 316–336. https://doi.org/10.1016/j. jsg.2014.05.010. Díaz-Azpiroz, Brune, S., Leever, K.A., Fernández, C., Czech, D.M., 2016. Tectonics of oblique plate boundary systems. Tectonophysics 693, 165–170. Díaz-Azpiroz, M., Fernández, C., Czeck, D.M., 2018. Are we studying deformed rocks in the right sections? Best practices in the kinematic analysis of 3D deformation zones. Journal of Structural Geology https://doi.org/10.1016/j.jsg.2018.03.005. Enlow, R.L., Koons, P.O., 1998. Critical wedges in three dimensions: analytical expressions from Mohr-Coulomb constrained perturbation analysis. Journal of Geophysical Research 103, 4897–4914. Escuder-Viruete, J., Baumgartner, P.O., 2014. Structural evolution and deformation kinematics of a subduction-related serpentinite-matrix mélange, Santa Elena peninsula, northwest Costa Rica. Journal of Structural Geology 66, 356–381. Escuder-Viruete, J., Valverde-Vaquero, P., Rojas-Agramonte, Y., Gabites, J., Pérez-Estaún, A., 2013. From intra-oceanic subduction to arc accretion and arc continent collision: insights from the structural evolution of the Río San Juan metamorphic complex, northern Hispaniola. Journal of Structural Geology 46, 34–56. Fernández, C., Díaz-Azpiroz, M., 2009. Triclinic transpression zones with inclined extrusion. Journal of Structural Geology 31, 1255–1269. Fernández, C., Czeck, D.M., Díaz Azpiroz, M., 2013. Testing the model of oblique transpression with oblique extrusion in two natural cases: steps and consequences. Journal of Structural Geology 54, 85–102. Festa, A., Pini, G.A., Dilek, Y., Codegone, G., 2010. Mélanges and mélange-forming processes: a historical overview and new concepts. In: Dilek, Y. (Ed.), Alpine Concept in Geology International Geology Review. 52(10–12), pp. 1040–1105. Festa, A., Dilek, Y., Pini, G.A., Codegone, G., Ogata, K., 2012. Mechanisms and processes of stratal disruption and mixing in the development of mélanges and broken formations: redefining and classifying mélanges. Tectonophysics 568, 7–24. Fitch, J.T., 1972. Plate convergence, transcurrent faults, and internal deformation adjacent to Southeast Asia and the Western Pacific. Journal of Geophysical Research 77, 4432–4460. Fossen, H., Tikoff, B., 1993. The deformation matrix for simultaneous simple shearing, pure shearing and volume change, and its application to transpression-transtension tectonics. Journal of Structural Geology 15, 413–422. Fossen, H., Tikoff, B., 1998. Extended models of transpression and transtension, and application to tectonic settings. In: Holdsworth, R.E., Strachan, R.A., Dewey, J.F. (Eds.), Continental Transpression and Transtension Tectonics. Geological Society, London, Special Publication 135, pp. 15–33. Fossen, H., Teyssier, C., Whitney, D., 2013. Transtensional folding. Journal of Structural Geology 56, 89–102. Fuentes, P., Díaz-Alvarado, J., Fernández, C., Díaz-Azpiroz, M., Rodríguez, N., 2016. Structural analysis and shape-preferred orientation determination of the mélange facies in the Chañaral mélange, Las Tórtolas Formation, Coastal Cordillera, northern Chile. Journal of South American Earth Sciences 67, 40–56. Fuentes, P., Díaz-Alvarado, J., Rodríguez, N., Fernandez, C., Breitkreuz, C., Contreras, A., 2017. Geochemistry, petrogenesis and tectonic significance of the volcanic rocks of

the Las Tortolas Formation, Coastal Cordillera, northern Chile. Journal of South American Earth Sciences https://doi.org/10.1016/j.jsames.2017.11.006. García-Sansegundo, J., Farias, P., Heredia, N., Gallastegui, G., Charrier, R., Rubio-Ordóñez, A., Cuesta, A., 2014. Structure of the Andean Palaeozoic basement in the Chilean coast at 31°30′ S: geodynamic evolution of a subduction margin. Journal of Iberian Geology 40 (2), 293–308. Ghosh, S.K., 1982. The problem of shearing along axial plane foliations. Journal of Structural Geology 4, 63–67. Glodny, J., Echtler, H., Collao, S., Ardiles, M., Burón, P., Figueroa, O., 2008. Differential Late Paleozoic active margin evolution in South-Central Chile (37° S-40° S) - the Lanalhue Fault Zone. Journal of South American Earth Sciences 26, 397–411. Godoy, E., Lara, P.L., 1998. Hojas Chañaral y Diego de Almagro, Región de Atacama. Servicio Nacional de Geología y Minería. Mapas Geólogicos no. 5-6, map escale 1: 100 000, Santiago. Grigull, S., Krohe, A., Moos, C., Wassmann, S., Stöckhert, B., 2012. “Order from chaos”: a field-based estimate on bulk rheology of tectonic mélanges formed in subduction zones. Tectonophysics 568, 86–101. Haq, S.S., Davis, D., 2010. Mechanics of forearc slivers: insights from simple analog models. Tectonics 29, TC5015. Heredia, N., García-Sansegundo, J., Gallastegui, G., Farias, P., Giacosa, R., Alonso, J.L., Busquets, P., Charrier, R., Clariana, P., Colombo, F., Cuesta, A., Gallastegui, J., Giambiagi, L., González-Menéndez, L., Limarino, C.O., Martín-González, F., Pedreira, D., Quintana, L., Rodríguez-Fernández, L.R., Rubio-Ordóñez, A., Seggiaro, R., Serra-Varela, S., Spalletti, L., Cardó, R., Ramos, V.A., 2016. Late Neoproterozoic-Paleozoic geodynamic evolution of the Andes of Argentina, Chile and the Antarctic Peninsula. Trabajos de Geologia 36, 237–278. Heredia, N., García-Sansegundo, J., Gallastegui, G., Farias, P., Giacosa, R., Hongn, F., Tubía, J.M., Alonso, J.L., Busquets, P., Charrier, R., Clariana, P., Colombo, F., Cuesta, A., Gallastegui, J., Giambiagi, L., González-Menéndez, L., Limarino, O., Martín-González, F., Pedreira, D., Quintana, L., Rodríguez-Fernández, L.R., Rubio-Ordóñez, A., Seggiaro, R., Serra-Varela, S., Spalletti, L., Cardó, R., Ramos, V.A., 2018. The Pre-Andean Phases of Construction of the Southern Andes Basement in Neoproterozoic–Paleozoic Times. In: Folguera, A., et al. (Eds.), The Evolution of the Chilean-Argentinean Andes, Springer Earth System Sciences, pp. 111–131 https://doi.org/10.1007/978-3319-67774-3_5. Hervé, F., 1988. Late Paleozoic subduction and accretion in Southern Chile. Episodes 11, 183–188. Hervé, F., Kawashita, K., Munizaga, F., Bassei, M., 1984. Rb-Sr isotopic ages from late Palaeozoic metamorphic rocks of Central Chile. Journal of the Geological Society (London) 141, 877–884. Hervé, F., Munizaga, F., Parada, M.A., Brook, M., Pankhurst, R.J., Snelling, N.J., Drake, R., 1988. Granitoids of the coast range of Central Chile: geochronology and geologic setting. Journal of South American Earth Sciences 1, 185–194. Hervé, F., Faundez, V., Calderón, M., Massone, H.-J., Willner, A.P., 2007. Metamorphic and plutonic basement complexes. In: Moreno, T., Gibbons, W. (Eds.), The Geology of Chile. Geological Society, London, pp. 5–19. Hervé, F., Calderón, M., Faning, C.M., Pankhurst, R.J., Godoy, E., 2013. Provenance variations in the Late Paleozoic accretionary complex of Central Chile as indicated by detrital zircons. Gondwana Research 23, 1122–1135. Hervé, F., Fanning, C.M., Calderón, M., Mpodozis, C., 2014. Early Permian to Late Triassic batholiths of the Chilean Frontal Cordillera (28°–31°S): SHRIMP U–Pb zircon ages and Lu–Hf and O isotope systematics. Lithos 184–187, 436–446. Hsü, K.J., 1968. Principles of mélanges and their bearing on the Franciscan-Knoxville Paradox. Geological Society of America Bulletin 79, 1063–1074. Hyppolito, T., Juliani, C., García-Casco, A., Meira, V.T., Bustamante, A., Hervé, F., 2014. The nature of the Palaeozoic oceanic basin at the southwestern margin of Gondwana and implications for the origin of the Chilenia terrane (Pichilemu region, Central Chile). International Geology Review 56 (9), 1097–1121. Hyppolito, T., Juliani, C., Garcia-Casco, A., Meira, V., Bustamante, A., Hall, C., 2015. LP/HT metamorphism as a temporal marker of change of deformation style within the Late Palaeozoic accretionary wedge of Central Chile. Journal of Metamorphic Geology 33 (9), 1003–1024. Jiang, D., 1994. Vorticity determination, distribution, partitioning and the heterogeneity and nonsteadiness of natural deformations. Journal of Structural Geology 16, 1–121. Jiang, D., 2007. Sustainable transpression: an examination of strain and kinematics in deforming zones with migrating boundaries. Journal of Structural Geology 29, 1984–2005. Jiang, D., Williams, P.F., 1998. High-strain zones: a unified model. Journal of Structural Geology 20, 1105–1120. Jiang, D., Lin, S., Williams, P.F., 2001. Deformation paths in high-strain zones, with reference to slip partitioning in transpressional plate-boundary regions. Journal of Structural Geology 23, 991–1005. Jones, R.J., Holdsworth, R.E., Bailey, W., 1997. Lateral extrusion in transpression zones: the importance of boundary conditions. Journal of Structural Geology 19, 1201–1217. Kamb, W.B., 1959. Ice petrofabric observations from Blue Glacier, Washington, in relation to theory and experiment. Journal of Geophysical Research 64, 1891–1909. Kano, K., Konishi, Y., 2001. Deformation process in the shallow levels of an accretionary prism: the Mesozoic Torlesse Terrane, South Island, New Zealand. The Island Arc 10, 158–174. Kato, T.T., Godoy, E., 2015. Middle to late Triassic mélange exhumation along a preAndean transpressional fault system: coastal Chile (26°-42° S). International Geology Review 57 (5–8), 606–628. Kimura, G., Kitamura, Y., Hashimoto, Y., Yamaguchi, A., Shibata, T., Ujiie, K., Okamoto, S., 2007. Transition of accretionary wedge structures around the up-dip limit of the seismogenic subduction zone. Earth and Planetary Science Letters 255, 471–484.

P. Fuentes et al. / Gondwana Research 74 (2019) 251–270 Kimura, G., Yamaguchi, A., Hojo, M., Kitamura, Y., Kameda, J., Ujiie, K., Hamada, Y., Hamhashi, M., Hina, S., 2012. Tectonic mélange as fault rock of subduction plate boundary. Tectonophysics 568-569, 25–38. Kitamura, Y., Kimura, G., 2012. Dynamic role of tectonic mélange during interseismic process of plate boundary mega earthquakes. Tectonophysics 568-569, 39–52. Kleiman, L.E., Japas, M.S., 2009. The Choiyoi volcanic province at 34°S-36°S (San Rafael, Mendoza, Argentina): implications for the Late Palaeozoic evolution of the southwestern margin of Gondwana. Tectonophysics 473, 283–299. Krohe, A., 2017. The Franciscan Complex (California, USA) – the model case for returnflow in a subduction channel put to the test. Gondwana Research 45, 282–307. Kusky, T.M., Bradley, D.C., 1999. Kinematic analysis of mélange fabrics: examples and applications from the McHugh Complex, Kenai Peninsula, Alaska. Journal of Structural Geology 21, 1773–1796. del Rey, A., Arriagada, C., Dekart, K., Martínez, F., 2016. Resolving the paradigm of the late Paleozoic-Triassic Chilean magmatism: isotopic approach. Gondwana Research 37, 172–181. Lash, G.G., 1987. Diverse mélanges of an ancient subduction complex. Geology 15, 652–655. Lawver, L.A., Scotese, C.R., 1987. A revised reconstruction of Gondwanaland. In: McKenzie, G.D. (Ed.), Gondwana Six: Structure, Tectonics, and Geophysics. American Geophysical Union Geophysical Monograph 40, pp. 17–23. Leever, K.A., Gabrielsen, R.H., Sokoutis, D., Willingshofer, E., 2011. The effect of convergence angle on the kinematic evolution of strain partitioning in transpressional brittle wedges: insight from analog modeling and high-resolution digital image analysis. Tectonics 30, TC2013. Lin, S., Jiang, D., Williams, P.F., 1998. Transpression (or transtension) zones of triclinic symmetry: natural example and theoretical modeling. In: Holdsworth, R.E., Strachan, R.A., Dewey, J.F. (Eds.), Continental Transpression and Transtension Tectonics. Geological Society vol. 135. Special Publication, London, pp. 41–57. Lister, G.S., Williams, P.F., 1983. The partitioning of deformation in flowing rock masses. Tectonophysics 92, 1–33. Marioth, R., 2001. Characterisierung und Quantifizierung thermischer und diagenetischer Prozesse im karbonischen Akkretionsprisma in Nordchile. (Unpublished PhD thesis). Univ. Heidelberg, Germany (144 pp.). Martin, M.W., Kato, T.T., Rodríguez, C., Godoy, E., Duhart, P., McDonough, M., Campos, A., 1999. Evolution of the Late Paleozoic accretionary complex and overlying forearcmagmatic arc, south Central Chile (38°-41° S): constraints for the tectonic setting along the southwestern margin of Gondwana. Tectonics 18, 582–605. McCaffrey, R., 1992. Oblique plate convergence, slip vectors and forearc deformation. Journal of Geophysical Research 97, 8905–8915. McCaffrey, R., Zwick, P.C., Bock, Y., Prawirodirdjo, L., Genrich, J.F., Stevens, C.W., Puntodewo, S.S.O., Subarya, C., 2000. Strain partitioning during oblique plate convergence in northern Sumatra: geodetic and seismologic constraint and numerical modeling. Journal of Geophysical Research 105, 28,363–28,376. McClay, K.R., Whitehouse, P.S., Dooley, T., Richards, M., 2004. 3D evolution of fold and thrust belts formed by oblique convergence. Marine and Petroleum Geology 21, 857–877. Meneghini, F., Marroni, M., Moore, J.C., Pandolfi, L., Rowe, C.D., 2009. The processes of underthrusting and underplating in the geologic record: structural diversity between the Franciscan Complex (California), the Kodiak Complex (Alaska) and the Internal Ligurian Units (Italy). Geological Journal 44, 126–152. Miller, H., 1970. Vergleichende Studien an prämesozoischen Gesteine Chile unter besonderen Berücksichtigung ihrer Kleintektonik. Geotektonische Forschungen 36, 1–64. Mpodozis, C., Kay, S.M., 1990. Provincias magmáticas ácidas y evolución tectónica de Gondwana: Andes chilenos (28-31°S). Andean Geology 17, 153–180. Mpodozis, C., Kay, S.M., 1992. Late Paleozoic to Triassic evolution of the pacific Gondwana margin: evidence from Chilean frontal Cordilleran batholiths. Geological Society of America Bulletin 104, 999–1014. Nabavi, S.T., Díaz-Azpiroz, M., Talbot, C.J., 2017. Inclined transpression in the Neka Valley, eastern Alborz, Iran. International Journal of Earth Sciences 106, 1815–1840. Naranjo, J.A., Puig, A., 1984. Hojas Taltal y Chañaral. Carta Geol. Chile. 62–63, pp. 1–140. Navarro, J.M., 2013. Petrotectónica del Complejo Metamórfico Punta de Choros, IIIIV Región, Chile. (Graduated thesis). University of Chile (110 pp.). Naylor, M., Sinclair, H.D., Willett, S., Cowie, P.A., 2005. A discrete element model for orogenesis and accretionary wedge growth. Journal of Geophysical Research 110, B12403. https://doi.org/10.1029/2003JB002940. Niwa, M., 2006. The structure and kinematics of an imbricate stack of oceanic rocks in the Jurassic accretionary complex of Central Japan: an oblique subduction model. Journal of Structural Geology 28 (9), 1670–1684. Ohsumi, T., Ogawa, Y., 2008. Vein structures, like ripple marks, are formed by shortwavelength shear waves. Journal of Structural Geology 30 (6), 719–724. Onishi, C.T., Kimura, G., 1995. Change in fabric of mélange in the Shimanto Belt, Japan: change in relative convergence? Tectonics 14, 1273–1289. Pankhurst, R.J., Hervé, F., Fanning, C.M., Calderón, M., Niemeyer, H., Griem-Klee, S., Soto, F., 2016. The pre-Mesozoic rocks of northern Chile: U-Pb ages, and Hf and O isotopes. Earth-Science Reviews 152, 88–105. Passchier, C.W., 1997. The fabric attractor. Journal of Structural Geology 19, 113–127. Pfiffner, O.A., Ramsay, J.G., 1982. Constraints on geological strain rates: arguments from finite strain rates of naturally deformed rocks. Journal of Geophysical Research 87 (B1), 311–321. Philippon, M., Corti, G., 2016. Obliquity along plate boundaries. Tectonophysics 693, 171–182. Pini, G.A., 1999. Tectonosomes and olistostromes in the argille scagliose of the northern Apennines, Italy. Geological Society of America Special Paper 335 (73 pp.).

269

Plunder, A., Thieulot, C., van Hinsbergen, D.J.J., 2018. The effect of obliquity on temperature in subduction zones: insights from 3-D numerical modeling. Solid Earth 9, 759–776. Ramos, V.A., 1988. The tectonics of the Central Andes; 30° to 33° S latitude. In: Clark, S., Burchfiel, D. (Eds.), Processes in Continental Lithospheric Deformation. Geological Society of America, Special Paper 218, pp. 31–54. Ramsay, J.G., 1967. Folding and Fracturing of Rocks. McGraw-Hill, New York (568 pp.). Ramsay, J.G., Graham, R.H., 1970. Strain variation in shear belts. Canadian Journal of Earth Sciences 7, 786–813. Ramsay, J.G., Huber, M.I., 1983. The Techniques of Modern Structural Geology. Volume 1: Strain Analysis. Academic Press, London (307 pp.). Ramsay, J.G., Huber, M.I., 1987. The Techniques of Modern Structural Geology. Volume 2: Folds and Fractures. Academic Press, London (393 pp.). Raymond, L.A., 1984. Classification of mélanges. In: Raymond, L.A. (Ed.), Mélanges: Their Nature, Origin and Significance. 198. Geological Society of America Special Papers, Boulder, Colorado, pp. 7–20. Raymond, L.A., Yurkovich, S.P., McKinney, M., 1989. Block-in-matrix structures in the North Carolina Blue Ridge belt and their significance for the tectonic history of the southern Appalachian orogen. In: Horton Jr., J.W., Rast, N. (Eds.), Mélanges and Olistostromes of the Appalachians. Geological Society of America Special Papers 228, pp. 195–216. Ring, U., Willner, A.P., Layer, P.W., Richter, P.P., 2012. Jurassic to Early Cretaceous postaccretional sinistral transpression in north-central Chile (latitudes 31-32° S). Geological Magazine 149, 208–220. Richter, P., Ring, U., Willner, A.P., Leiss, B., 2007. Structural contacts in subduction complexes and their tectonic significance: The Late Paleozoic coastal accretionary wedge of central Chile. Journal of the Geological Society of London 164, 203–214. Robin, P.Y.F., Cruden, A.R., 1994. Strain and vorticity patterns in ideally ductile transpressional zones. Journal of Structural Geology 16, 447–466. Sanderson, D.J., Marchini, W.R.D., 1984. Transpression. Journal of Structural Geology 6, 449–458. Scheuber, E., González, G., 1999. Tectonics of the Jurassic-Early Cretaceous magmatic arc of the north Chilean Coastal Cordillera (22°-26°S): a story of crustal deformation along a convergent plate boundary. Tectonics 18, 895–910. Schreurs, G., Colletta, B., 1998. Analogue modelling of faulting in zones of continental transpression and transtension. In: Holdsworth, R.E., Strachan, R.A., Dewey, J.F. (Eds.), Continental Transpressional and Transtensional Tectonics. Geological Society, London, pp. 59–79. Schulmann, K., Thompson, A.B., Lexa, O., Ježek, J., 2003. Strain distribution and fabric development modeled in active and ancient transpressive zones. Journal of Geophysical Research - Solid Earth 108 (B1). Shi, Y., Yu, J.H., Santosh, M., 2013. Tectonic evolution of the Qinling orogenic belt, Central China: new evidence from geochemical, zircon UePb geochronology and Hf isotopes. Precambrian Research 231, 19–60. Shreve, R.L., Cloos, H., 1986. Dynamics of sediment subduction, mélange formation, and prism accretion. Journal of Geophysical Research 91, 10229–10245. Sullivan, W.A., 2008. Significance of transport-parallel strain variations in part of the Raft River shear zone, Raft River Mountains, Utah, USA. Journal of Structural Geology 30, 138–158. Sullivan, W.A., Law, R.D., 2007. Deformation path partitioning within the transpressional White Mountain shear zone, California and Nevada. Journal of Structural Geology 29, 583–598. Suzuki, T., 1986. Melange Problem of Convergent Plate Margins in the Circum-Pacific Regions. Memoirs of the Faculty of Science, Kochi University, Series E. Geology 7 pp. 23–48. Taira, A., Byrne, T., Ashi, J., 1992. Photographic Atlas of an Accretionary Prism: Geological Structures of the Shimanto Belt. Springer-Verlag and University of Tokyo Press, Japan, Tokyo (124 pp.). Tikoff, B., Teyssier, C., 1994. Strain modeling of displacement-field partitioning in transpressional orogens. Journal of Structural Geology 16, 1575–1588. Truesdell, C.A., 1954. The Kinematic of Vorticity. Indiana University Press, Bloomington. Truesdell, C.A., Toupin, R.A., 1960. The classic field theory. In: Flügge, S. (Ed.), Encyclopedia of Physics. Volume III: Principles of Classical Mechanics and Field Theory. Springer-Verlag, Berlin, pp. 226–793. Ujiie, K., 2002. Evolution and kinematics of an ancient decollement zone, mélange in the Shimanto accretionary complex of Okinawa Island, Ryukyu Arc. Journal of Structural Geology 24, 937–952. Ulricksen, C., 1979. Regional Geology, Geochronology and Metallogeny of the Coastal Cordillera of Chile Between 25°30′ and 26°00′ South. (Thesis). Dalhousie University, Canada. Vannucchi, P., Bettelli, G., 2002. Mechanism of subduction accretion as implied from the broken formations in the Apennines, Italy. Geology 30, 835–838. Williams, P.F., 1976. Relationships between axial-plane foliations and strain. Tectonophysics 30, 181–196. Willner, A.P., 2005. Pressure-Temperature Evolution of a Late Paleozoic paired Metamorphic Belt in North-Central Chile (34°-35° 30′S). Journal of Petrology 46, 1805–1833. Willner, A.P., Pawlig, S., Massonne, H.-J., Hervé, F., 2001. Metamorphic evolution of spessartine quartzites (coticules) in the high pressure/low temperature complex at Bahia Mansa (Coastal Cordillera of Southern Central Chile). Canadian Mineralogist 39, 1547–1569. Willner, A.P., Glodny, J., Gerya, T.V., Godoy, E., Massonne, H.-J., 2004. A counterclockwise PTt-path in high pressure-low temperature rocks from the Coastal Cordillera accretionary complex of South Central Chile: constraints for the earliest stage of subduction mass flow. Lithos 73, 283–310. Willner, A.P., Thomson, S.N., Kröner, A., Wartho, J.A., Wijbrans, J.R., Hervé, F., 2005. Time markers for the evolution and exhumation history of a Late Paleozoic paired

270

P. Fuentes et al. / Gondwana Research 74 (2019) 251–270

metamorphic belt in North-Central Chile (34°-35° 30′ S). Journal of Petrology 46, 1835–1858. Willner, A.P., Gerdes, A., Massonne, H.J., 2008. History of crustal growth and recycling at the Pacific convergent margin of South America at latitudes 29-36 ºS revealed by a U-Pb and Lu-Hf isotope study of detrital zircon from late Paleozoic accretionary systems. Chemical Geology 253, 114–129. Willner, A.P., Massonne, H.-J., Ring, U., Sudo, M., Thomson, S.N., 2012. P-T evolution and timing of a late Palaeozoic fore-arc system and its heterogeneous Mesozoic overprint in north-central Chile (latitudes 31-32°S). Geological Magazine 149, 177–207. Yamamoto, Y., Ohta, Y., Ogawa, Y., 2000. Implication for the two-stage layer-parallel faults in the context of the Izu forearc collision zone: examples from the Miura accretionary prism, Central Japan. Tectonophysics 325 (1–2), 133–144.

Yamamoto, Y., Ogawa, Y., Uchino, T., Muraoka, S., Chiba, T., 2007. Large-scale chaotically mixed sedimentary body within the late Pliocene to Pleistocene Chikura group, Central Japan. Island Arc 16, 505–507. Yamamoto, Y., Nidaira, M., Ohta, Y., Ogawa, Y., 2009. Formation of chaotic rock units during primary accretion processes: examples from the Miura-Boso accretionary complex, central Japan. Island Arc 18, 496–512. Zheng, Y.F., Zhao, Z.F., 2017. Introduction to the structures and processes of subduction zones. Journal of Asian Earth Sciences 145, 1–15. Zhu, S., Cai, Y., Shi, Y., 2006. The contemporary strain rate field of continental China predicted from GPS measurements and its geodynamic implications. Pure and Applied Geophysics 163, 1477–1493.