Strike-Slip Shear Zones of the Iberian Massif: Are They Coeval?

Strike-slip shear zones of the Iberian Massif: Are they coeval?

Rubén Díez Fernández1,2,* and Manuel Francisco Pereira2 1DEPARTAMENTO DE GEODINÁMICA, FACULTAD DE CIENCIAS GEOLÓGICAS, UNIVERSIDAD COMPLUTENSE DE MADRID, C/ JOSÉ ANTONIO NOVAIS 2, 28040 MADRID, SPAIN 2INSTITUTO DOM LUIZ, DEPARTAMENTO DE GEOCIÊNCIAS, ESCOLA DE CIÊNCIAS E TECNOLOGIA, UNIVERSIDADE DE ÉVORA, APARTADO 94, 7001-554 ÉVORA, PORTUGAL

ABSTRACT

Strike-slip shear zones of the Variscan orogen are used to derive the evolution of paleostrain and discuss the kinematics of the waning stages of the Gondwana-Laurussia collision during the amalgamation of Pangea. In the Iberian Massif, the recognition of three late Carboniferous

deformation events related to strike-slip tectonics (D3, D4, D5) in the Trancoso-Pinhel region (Portugal) reveals that late orogenic transcurrent deformation was episodic and occurred in a short period of time (<15 m.y.). Early stages of strike-slip deformation included dextral and sinistral

shear zones and orogen-parallel upright folds (D3; ca. 311 Ma). These structures followed the development of extensional shear zones (D2) dur-

ing the tectonothermal reequilibration of the orogen. D3 structures were deflected and folded by the sinistral 4D Juzbado-Penalva do Castelo shear zone, dated as ca. 309–305 Ma by SHRIMP (sensitive high-resolution ion microprobe) U-Pb zircon dating of synkinematic granitoids.

D3 and D4 structures were folded under east-west compression (D5) influenced by the strike-slip movement of the dextral Porto-Tomar shear

zone. Variscan movement along the Porto-Tomar shear zone started ca. 304 Ma (onset of the Buçaco basin and syn-D5 granites), but ceased before ca. 295 Ma (age of the final closure of the Ibero-Armorican arc and crosscutting granites). The contrasting geometry, kinematics, and timing of these strike-slip shear zones are explained by deformation partitioning upon a rheologically inhomogeneous crust with structural and tectonothermal anisotropies generated during previous deformation. The convergence vector between Gondwana and Laurussia during

D3–D5 remained the same, and was equivalent to the vector that explains the previous tectonic record (D2) in central and northwestern Iberia.

LITHOSPHERE; v. 9; no. 5; p. 726–744; GSA Data Repository Item 2017250 | Published online 30 June 2017 https://doi.org/10.1130/L648.1

INTRODUCTION Strike-slip shear zones are common in orogenic systems, and are considered one of the tools to determine relative plate movements (e.g., Structures with contrasting geometry and kinematics can be (1) the Shelley and Bossière, 2002). Block extrusion or escape tectonics may result of a single phase of deformation operating differently from one produce conjugated faults (e.g., Tapponnier and Molnar, 1979; Tappon- place to the other due to, e.g., local variations of finite strain, strain vor- nier et al., 1982), so using a suitable general picture, including timing and ticity, thermal conditions, rheological rock properties, and presence of fault trace, is advised for producing tectonic models based on the study of fluids; (2) independent from each other and derive from different phases strike-slip shear zones. Here we provide a case study of the complexity of deformation; or (3) the result of progressive deformation associated regarding the structural evolution, kinematic interpretation at the scale of with gradual changes of pressure-temperature conditions, e.g., emplace- the orogen, and dating of strike-slip shear zones formed in a collisional ment and cooling of a syntectonic granitoid (e.g., Carreras et al., 2004). orogen. In the Variscan orogen, some major geotectonic domains are Distinguishing between these three scenarios is challenging when the bounded by strike-slip shear zones (Fig. 1). Defining their kinematics and structures under consideration share geometrical properties, and/or when timing beyond the limits of analytical methods is thus necessary to better the time period for their development is shorter than the time resolution understand the processes involved in the building of long-lived orogenic of absolute dating methods. Establishing the geometry and timing for systems such as the Variscan. The strike-slip shear zones analyzed here structures formed during orogeny is key to reconstructing the associated formed at an advanced stage in the late Paleozoic collision of Gondwana convergence and/or divergence and evolution. Inferences about evolving and Laurussia, and therefore provide a closer view into the kinematics paleostrain can help to build tectonic models for orogens where significant and finite strain during the final amalgamation of Pangea. parts of the paleogeographic references are modified, such as in the case In the Iberian Massif of the Variscan orogenic system, strike-slip shear of Paleozoic orogens that are now dispersed over different tectonic plates. zones and upright folds are considered the result of deformation partition- Therefore, identifying the full sequence of structures in a given region is ing in a transpressional regime. However, the timing for transpression has a fundamental step toward reconstructions of large-scale tectonic settings. been assigned to contrasting periods in the Variscan orogeny (e.g., Iglesias Ponce de Leon and Choukroune, 1980; Dias and Ribeiro, 1998; Dias et al., 2010, 2013; Díez Fernández and Martínez Catalán, 2012; Martínez Ruben Díez Fernández http://orcid.org/0000 -0002 -0379 -7970 Catalán, 2012). The debate is centered on orogen-parallel upright folds *Corresponding author: [email protected] accompanying the strike-slip shear zones. Were they developed during the

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EUROPEAN VARISCIDES the sequence of geological processes considered is fast enough not to be (CROPPING OUT / COVERED) fully tracked by isotopic dating methods. Here we present field structural IAN AR M C HE data from an area of the Iberian Massif selected because of its potential AVALONIAN FORELAND THRUST BELT BO (RHENOHERCYNIAN ZONE) BOHEMIAN to examine the relationships between the late strike-slip shear zones of RHEIC SUTURE MASSIF the Variscan orogen in the Trancoso-Pinhel region. Our data reveal the (REWORKED) existence of three different pulses of strike-slip deformation in the Central INTERNAL VARISCAN ZONE WITH ALLOCHTHONOUS UNITS Iberian Zone of the Iberian Massif between ca. 311 Ma and ca. 300 Ma. Structural data are reinforced by SHRIMP (sensitive high-resolution ion PARAUTOCHTHON microprobe) U-Pb zircon dating of a synkinematic granitoid. We use our AUTOCHTHON structural and geochronological data to demonstrate the diachronous and FORELAND stepped character of the late Carboniferous strike-slip shear zones of the THRUST BELT Iberian Massif, and constrain the kinematic framework of the waning RHENISMASSIFH GONDWANA stages of the Variscan orogeny.

GEOLOGICAL SETTING MASSIF CENTRAL The Iberian Massif is located in the southern branch of the Variscan orogen (Fig. 1), which is conventionally attributed to the progressive collision between Gondwana and Laurussia in the late Paleozoic (Matte, 2001; Faure et al., 2009; Martínez Catalán et al., 2009; Arenas et al., 2014; Díez

C Fernández et al., 2016). The Central Iberian Zone represents a section of R MASSIF A ARMORICAN ? L the margin of Gondwana underlying a set of far-traveled allochthonous ? A MA R SSI T F CE N terranes that spread across the Iberian Massif (Fig. 1; Díez Fernández ORICA RM N -A 0 200 km O RC and Arenas, 2015). The allochthonous terranes are tectonic slices of con- R A

B I tinental and oceanic crust that were originally located outboard of and 7 along the margin of Gondwana before collision (e.g., Arenas et al., 2016), C E 1 N T during which they were transported onto inner sections of the margin 2 R A L

I 4 5 3 B (e.g., Martínez Catalán et al., 2009; Díez Fernández et al., 2016). Late FIGURE 2 E

6 RI

A MAIN VARISCAN Carboniferous strike-slip shear zones (Fig. 2) cut across folds and crustal-

A STRIKE-SLIP

R scale thrusts formed during the early stages of collision and the emplace-

1 IBERIAN C SHEAR ZONES ment of allochthonous units, and are spatially and temporally related to 8 MASSIF RHEIC SUTURE (REWORKED) extensional shear zones and elongated granitic massifs. Accompanying 9 MAIN VARISCAN the strike-slip shear zones, late upright folds affect the thrust nappes, the Central THRUSTS extensional shear zones, and the granitoids (Fig. 3). Iberian The Trancoso-Pinhel region (Fig. 2) is located in the southern branch Zone ALPINE FRONT of the Ibero-Armorican arc and near the broad axial zone of the Central Figure 1. Location of the study area in the Variscan orogen (after Díez Fernán- Iberian arc (Fig. 1). The section consists of Ediacaran to Ordovician dez and Arenas, 2015). Note the location of the Central Iberian Zone. Main metasedimentary rocks that alternate with rare bodies of orthogneiss, Variscan strike-slip shear zones of the Iberian Massif: 1—Porto-Tomar; 2—Mal- pica-Lamego; 3—Juzbado-Penalva do Castelo; 4—Douro-Beira; 5—Huebra; and are intruded by abundant and variably deformed ca. 331–304 Ma 6—Tamames; 7—Palas de Rei; 8—Coimbra-Córdoba; 9—South Iberian. granitoids (Díez Fernández and Pereira, 2016, and references therein) (Figs. 4A, 4B). Variscan metamorphism ranges between greenschist and granulite facies conditions (Regêncio Macedo, 1988; Valle Aguado et al., early stages of collision (e.g., Dias et al., 2010, 2013), or at an advanced 1993; Pereira, 2014) and the grade of regional metamorphism increases stage (e.g., Martínez Catalán, 2012)? These opposite views are not mutu- downsection. The regional structure is defined by an upright synform, ally exclusive, because detailed structural studies have consistently demon- cored by Ordovician quartzites and black phyllites (Marofa synform), strated the existence and interference of earlier and later upright folds (Díez that continues for tens of kilometers to the southeast and northwest (Fig. Balda, 1986; Díez Balda et al., 1990a, 1990b, 1995; Díez Balda and Vegas, 2; Tamames-Marofa-Sátão synform). Strike-slip shear zones, such as the 1992; Díez Fernández et al., 2013; Díez Fernández and Pereira, 2016). Huebra and the Juzbado-Penalva do Castelo shear zones, are important Even if a late orogenic nature is admitted for the strike-slip shear zones local structures (Figs. 2, 4A, and 4B). of the Iberian Massif, an increasing amount of geochronological data suggests that their formation took place over a short time interval, from ca. DEFORMATION PHASES 315 to 300 Ma (Regêncio Macedo, 1988; Rodríguez et al., 2003; Valle Aguado et al., 2005; Gutiérrez-Alonso et al., 2015; Díez Fernández and The Trancoso-Pinhel region is affected by five main phases of deforma-

Pereira, 2016). Consequently, models considering all of the strike-slip tion (D1–D5), and by a set of late faults (Díez Fernández and Pereira, 2016).

shear zones as coeval have emerged (e.g., Gutiérrez-Alonso et al., 2015). Upright folds characterized by an axial planar foliation (S1) represent D1.

Isotopic dating of tectonic fabrics is a tool to determine the evolu- D2 records development of the Pinhel shear zone, a low-dipping, ductile, tion of orogenic systems. However, the principle of superposition, when extensional shear zone with a penetrative medium- to high-grade foliation

applied to the recognition of the sequence of phases of deformation using (S2) that overprints D1 structures. D2 was responsible for telescoping of field observations, is not rivaled by any radiometric dating method in regional metamorphic isograds, exhumation of migmatized rocks, and the

terms of relative timing. Such observations are particularly useful when progressive deformation of synkinematic granitoids. D2 metamorphism

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0 50 km

42º

MLSZ DBSZ A

FIG. 4A

PTSZ VS

Z 5 TMSS 5 JPSZ HSZ 41º

DBSZ 3 2 4 2 JPSZ TMSS 3 1 B 4 C 1 TSZ

FIG. 4B

Buçaco basin

PTSZ D

40º E

COVER 8º AUTOCHTHON 7º VARISCAN MAGMATISM 6º TRIASSIC - QUATERNARY ORDOVICIAN - DEVONIAN LATE KINEMATIC GRANODIORITES - GRANITES ALLOCHTHONOUS COMPLEXES ORDOVICIAN (Armorican Quartzite includ.) POST-KINEMATIC GRANODIORITES - GRANITES UPPER UNITS CAMBRIAN-ORDOVICIAN MAGMATISM POST KINEMATIC OPHIOLITIC UNITS EDIACARAN - CAMBRIAN CORDIERITE GRANITOIDS SYN-KINEMATIC BASAL UNITS EDIACARAN TWO-MICA LEUCOGRANITES SYN-KINEMATIC SYNOROGENIC DEPOSITS PRE-VARISCAN ORTHOGNEISSES GRANODIORITES - MONZOGRANITES DEVONIAN - CARBONIFEROUS D ANTIFORM 3 LOW-ANGLE THRUSTS

PARAUTOCHTHON mostly Central Iberian Zone D3 SYNFORM HIGH-ANGLE FAULTS EDIACARAN? - DEVONIAN 3 Granitoid massifs 3 LOCALITIES STRIKE-SLIP SHEAR ZONE

Figure 2. Regional map of the central part of the Iberian Massif (after Díez Fernández and Pereira, 2016). The traces of major faults, late upright folds, and strike-slip shear zones are shown. Granitoid massifs (numbered white circles): 1—Cota-Viseu (ca. 306 Ma; Costa, 2011); 2—Aguiar da Beira (ca. 304 Ma; Costa, 2011); 3—Mêda-Escalhão (ca. 317 Ma; Costa, 2011); 4—Villavieja de Yeltes (ca. 304 Ma; Gutiérrez-Alonso et al., 2011); 5—Lavadores. Localities (numbered black circles): 1—Tamames; 2—Marofa (mountain peak); 3—Sátão; 4—Penalva do Castelo; 5—Juzbado. Abbreviations: DBSZ—Douro-Beira shear zone; HSZ—Huebra shear zone; JPSZ—Juzbado-Penalva do Castelo shear zone; MLSZ—Malpica-Lamego shear zone; PTSZ—Porto-Tomar shear zone; TMSS—Tamames-Marofa-Sátão synform; TSZ—Tamames shear zone; VSZ—Villalcampo shear zone.

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Juzbado-Penalva do Castelo Huebra shear zone shear zone Tamames-Marofa-Sátão synform

015 km Morais complex

Mêda-Escalhão Pinhel shear zone massif SECTION A - B SECTION C - D SECTION D - E NNE SSW NNE SSW NE SW

Figure 3. Synthetic geological section across the southern part of the Central Iberian Zone (after Díez Fernández and Pereira, 2016). The legend and trace of the composite cross section are indicated in Figure 2.

created two contrasting metamorphic domains, a low-grade domain (LGD) domain, and the domain located away from the core of the shear zone characterized by greenschist facies metamorphic rocks that occupy the have similar trends (see following for details) and exhibit similar varia- hanging wall of the Pinhel shear zone, and a high-grade domain (HGD) tions in plunge, but they do not interfere with each other, suggesting that

featuring amphibolite to granulite facies metamorphic rocks that occupy they also developed during D3.

the footwall (Figs. 3 and 4; Díez Fernández and Pereira, 2016). In the high-strain domain, S3 is accompanied by the shallowly plung-

The third phase of deformation (D3) produced widespread, shallowly ing L3e, and displays sinistral asymmetric fabrics (e.g., sigma objects, plunging upright folds along with a penetrative, low-grade axial planar C’, C’-S and S-C structures). Discrete strike-slip shear zones (decimeter

foliation (S3). The plunge of D3 folds varies from the east and west, produc- scale) are dispersed to the north of the core of the Huebra shear zone (not ing the pinch and swell map pattern that characterizes the region. To the represented in the map [Figs. 4A and 5A]). The lateral continuity of these

south of the Marofa synform, the Almeida-Malpartida granitic massif is shear zones could not be established due to limited exposure. D3 fold an intrusion that is discordant with respect to its previously deformed and axial traces located to the north of the Huebra shear zone show a fan-like metamorphosed host rocks (Fig. 4B). The N50°–60°E trending boundary pattern converging to the southeast in the core of the Huebra shear zone, of this nonfoliated biotite-rich granitic massif cuts across the LGD-HGD thus providing regional-scale criteria for sinistral kinematics (Fig. 5C).

boundary (i.e., Pinhel shear zone) and the axial traces of D3 folds. Regional mapping of the N80°–90°E trending Huebra shear zone sug-

D3 is also characterized by strike-slip shear zones that, together with gests that it may merge with the dextral, N130°–140°E trending Malpica-

D4 and D5, are described in the following. Lamego shear zone. Both shear zones are parallel to the axial planes of

major D3 folds (Fig. 2).

Huebra and Tamames Shear Zones (D3) The Tamames shear zone (Díez Fernández and Pereira, 2016) was described by Díez Balda et al. (1990a), who identified a strike-slip shear The Huebra shear zone occurs along the northern boundary of the zone trending N130°–140°E along the northeastern limb of the upright Mêda-Escalhão massif (Figs. 2 and 4). This shear zone shows subverti- fold that was named the Tamames syncline. Orthogonal shortening associ-

cal foliation (S3; Fig. 4A) and subhorizontal elongation lineation (L3e; ated with the strike-slip shear zone produces shallowly plunging upright

Fig. 5A). The strike of foliation (Fig. 4E) and trend of lineation (Fig. 5D) folds (attributed to the third phase of deformation) that interfere with D1

are parallel to the trace of the shear zone. Foliation planes show slight folds and produce minor domes and basins. A Z-shaped deflection of D1 variations in dip direction (to the north and to the south) and in inclina- folds in that area suggests a right-lateral movement for the Tamames shear tion values. The plunge of lineations also varies, and may be gently to zone (Díez Balda et al., 1990a). the east, horizontal, or gently to the west. The relative timing between the Huebra and Tamames shear zones Strain in the Huebra shear zone is heterogeneous. Deformation con- could not be established directly. They are separated by another strike-slip centrates along the northern boundary of the Mêda-Escalhão massif, and shear zone (Fig. 2; Juzbado-Penalva do Castelo shear zone; see following), diminishes progressively to the north and to the south. The deformation which impedes the use of in situ field criteria to address this question.

gradient is steeper to the south, where the Mêda-Escalhão massif is bound However, because both shear zones formed together with D3 folds, this by a narrow band of mylonites with subvertical foliation (Fig. 6A; for color particular set of strike-slip shear zones can be considered coeval. versions, see the GSA Data Repository Item1), which grades into a non-

penetrative, subvertical planar fabric defined by the preferred orientation Juzbado-Penalva do Castelo Shear Zone (D4) of feldspar phenocrysts. To the north of the shear zone, the metasedimentary rocks show more spaced crenulation cleavage and increasingly open The Juzbado-Penalva do Castelo shear zone (JPSZ) is a strike-slip upright folds (Fig. 6B) compared to those that characterize the high-strain system that is parallel to the northern limb of the Tamames-Marofa-Sátão

core (Fig. 6C). Fold axes (F3) in the domains with higher strain coincides D3 synform in the Trancoso-Pinhel region (Fig. 4B). Mapping to the east

with L3e (Fig. 6D). Upright folds of the high-strain domain, the transitional of these two major structures indicates that the JPSZ cuts and deflects

this D3 fold in a sinistral sense, affecting its trace for more than 100 km (Fig. 2). JPSZ represents the fourth phase of deformation in the study 1 GSA Data Repository Item 2017250, containing field photographs and Table DR1: U-Pb geochronology data of samples TP-4 and TP-13, is available at http://www area (D4) and separates the D3 Huebra shear zone in the north from the

.geosociety.org/datarepository/2017, or on request from [email protected]. D3 Tamames shear zone in the south (Fig. 2).

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Downloaded from http://pubs.geoscienceworld.org/gsa/lithosphere/article-pdf/9/5/726/3712177/726.pdf by guest on 29 September 2021 DÍEZ FERNÁNDEZ AND PEREIRA 60 40 50 30 45 45 26 45 22 30 80 50 36 58 68 10 0 55 ES 65 N 45 22 30 6 10 60 20 16 680000 680000 80 28 75 75 70 452000 50 45 50 18 35 70 70 S 60 68 22 60 40 20 55 14 35 40 30 70 86 40 50 55 km 70 SOUT H 65 40 45 50 40 57 50 80 60 A 24 15 84 32 17 45 32 30 80 50 39 45 53 70 70 ES 60 NE S 75 -CONGLOMER AT LOCALITIE NORT H 75 60 60 60 65 52 88 45 50 23 65 28 32 53 46 70 42 1 85 02 Mêda-Escalhão Massi f 26 30 60 36 58 48 70 30 50 20 45 78 72 51 OIDS (ca. 331-321 Ma ) 50 60 76 20 20 OIDS (ca. 321-317 Ma ) AND SCHISTS / BLACK SCHIST S 80 81 70 -SANDS TO 80 70 80 80 65 60 NES, ME TA 70 81 60 44 20 70 60 45 80 36 C 52 78 78 52 A’ 85 A-CONGLOMER AT 40 62 38 29 75 80 AND PHYLLITE S 72 80 680000 Almeida-Malpartida Massif 58 70 40 80 B’ AND ME TA 70 85 50 TISM 21 S -SANDS TO 30 25 80 57 60 NES 35 75 72 55 59 78 TA ARAGNEISSES 84 67 AND MET TO 60 50 70 85 80 50 65 62 45 81 86 70 Y DEFORMED GRANIT 55 45 40 Y DEFORMED GRANIT 51 40 56 63 45 70

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t 52 62 40 70 80 34 52 44 ) 30 50 40 80 70 68 24 70 55 75 60 45 40 65 regional metamorphism) 20 40 ) 55 78 70 71 2 60 Sill ou GRANITES (317 Ma ) 80 DEFORMED TWO-MICA 85 65 70 74 70 54 62 regional metamorphism) 60 45 50 46 80 34 2 70 40 45 80 30 70 40 75 84 85 56 46 30 24 20 80 60 45 3 RISCAN MAGMA 50 RIAB LY 38 36 60

Sill in 70 53 70 85 46 50 12 37 70 80 15 55 60 45 60 TWO-MICA VA MUSCOVITE GRANITES (305 Ma ) 50 80 t 70 75 40 80 67 54 VA RISCAN MAGMA RISCAN MAGMA 75 53 3 80 55 35 40 40 40 53 70 40 80 Bt ou 69 Contact (lithological Fault (high-angle LGD-HGD boundary (Pinhel shear zone Biotite in (M Juzbado-Penalva do Castelo faul Sillimanite in (M 60 -D VA VA Bt in 50 50 70 60 55 34 2 3 4 57 45 50 81 80 20 80 45 50 45 70 D D D 74 70 70 60 71 45 60 73 55 70 45 A’ 80 80 80 40 660000 60 4545000 72 70 70 A 56 75 76 65 80 64 61 85 66 (180°/67°) 55 70 46 50 52 83 45 70 Santa Eufémia Massif 70 64 TIS M 65 25 75 60 68 55 4530000 80 36 80 72 55 85 32 67 75 80 72 55 85 38 (poles to planes) SOUT H 80 ) 61 63 45 3 56 80 17 77 75 30 70 70 68 B’ 44 38 56 80 80 MEAN PRINC. ORIENT . 75 65 44 40 68 80 42 85 80 75 38 2 50 52 63 S 60 55 RISCAN MAGMA 60 60 85 65 10 70 3 85 40 FOLIATION (S S 82 75 50 80 80 3 79 ieiro Massi f VA 70 80 PORPHYRITIC MONZONITIC GRANITES (304 Ma ) MONZONITIC GRANITES QUARTZ DIKE D (n=919) 4 52 50 30 70 76 80 2 80 70 84 43 75 32 35 55 22 70 70 65 80 85 Post- D 80 77 45 São Pedro-V 50 54 40 74 60 80 3 50 50 in the northern part of the Trancoso-Pinhel region (lower hemisphere, equal angle). Universal Transverse Mercator coordinates coordinates Mercator Transverse Universal equal angle). hemisphere, (lower region Trancoso-Pinhel of the part in the northern 65 S 62 68 3 ° 77 70 70 80 65 74 64 57 78 70 57 70 80 80 65 80 75 ? 650000 ? 72 55 70 70 60 F 50 85 ROSE (n=919) Sector angle = 10° Maximum = 19,8% [182 data] Mean = 91°-271 70 65 72 85 80 74 85 70 55 70 3 80 74 S 80 66 74 78 74 83 (21°/87°) 70 30 44 31 77 25 E 70 Santa Eufémia Massi f 80 57 2 65 44 15 S (poles to planes) 80 60 ) 55 75 3 60 Pinhel 76 4510000 58 shear zone 36 4 65 MEAN PRINC. ORIENT . S 60 75 3 S 65 ) kinematic s ) 0 60 ) ) ) 40 2 57 4 30 68 30 D (top-to-the-SE) S 34 73 60 shear zone 55 44 FOLIATION (S 60 3 Juzbado-Penalv a 60 80 45 86 50 D (n=130) 68 44 28 3 17 S 55 40 f 45 30 44 84 n n n n 20 48 80 40 70 62 80 ° 640000 60 Axial trace (Ant-Sync Axial trace (Ant-Synf Axial trace (Ant-Synf Axial trace (Ant-Synf 65 4520000

Foliatio Foliatio Foliatio Foliatio 2 65 S 3 2 4 5 1 2 3 4 26 30 1 NORT H Sedimentary layering (S S S S S D D D D 65 contour lines 72 70 D 50 50 50 1% 2% 4% 8% 16 % 50 50 B ROSE (n=130) Sector angle = 10° Maximum = 37,7% [49 data] Mean = 111°-291 B Aguiar da Beira Massi Figure 4. (A) Geological map of the northern part of the Trancoso-Pinhel region including the axial traces of folds, the main planar features and the trace of selected metamorphic metamorphic of selected and the trace the main planar features of folds, including the axial traces region Trancoso-Pinhel of the part (A) Geological map of the northern 4. Figure region. Trancoso-Pinhel the of part southern the of map Geological (B) domain. HGD—high-grade domain; LGD—low-grade 2. Figure in indicated is location regional The isograds. 2—Santa 1—Trancoso; circles): black Localities (numbered region). (Trancoso-Pinhel area of the study structure and faulted the folded sections showing D) Geological cross (C, (E) synf—synform. sync—syncline; Ant—antiform; 8—Escalhão (not in the map). 7—Almendra; 6—La Fregeneda; Rodrigo; de Castelo 5—Figueira 4—Marofa; 3—Pinhel; Eufémia; and dip of S the trend plot showing Stereographic and dip of S the trend plot showing (F) Stereographic orientation. Orient.—principal Princ. 29). (zone

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52 l il

Sill Sill

4 ou S 5 ] 670000 e) 3 i) ) 3 3 7 3 Bt in Bt ou t 25 14 (274°/2° ) (258°/13° ) (268°/6° ) 10 18 m) 17 22 3 ° 20 20 (95°/5° ) 26 ° SYMBOLS 15 7 5 20 30 25 14 20 9 7 10 14 elongation lineation (L crenulation lineation (F crenulation / stretching lineation (F crenulation lineation (F elongation lineation (inclin./horiz.) (L intersection lineation (L crenulation lineation (F 15 10 26 10 4 5 2 3 3 3 4 8 15 16 3 5 7 D D D D D D D 7 MEAN PRINC. DIR. MEAN PRINC. DIR. MEAN PRINC. DIR. 12 10 11 5 5 10 11 5 6 6 8 MEAN PRINC. DIR. 20 52 4 30 17 ROSE (n=315) Sector angle = 10 Maximum = 16,8% [53 data Mean = 91°-271 10 6 7 18 6 6 20 6 6 8 STRUCTURAL 10 25 ) ) 15 INTERSECT. LINEATION (L 10 ELONGATION LINEATION (L 0 22 3 12 28 CRENULATION LINEATION (F MINERAL LINEATION (L -S D (n=23) 30 25 3 3 3 12 670000 D (n=144 D S (n=141 (n=7 ) 13 5 22 18 8 18 3 8 8 contour lines 30 8 22 21 20 7 10 24 37 10

20 21 l in l 45 16 17 1% 2% 4% 8% 16%

15 S l Si 35 11 17 50 8 5 12 15 25 25 17 8 13 14 4 15 660000 20 16 14 12 4 20 9 SOUT H 7 4 14 15 14 15 40 10 32 8 14 12 8 660000 22 30 5 15 7 18 28 7

t 23 22 20 5 13 54 51 10 3 km 20 16 6 5 17

Sill ou NORT H 20 15 35 10 20 19 ) 45 30 40 43 12 20 10 6 20 25 32 6 8 5 23 5 5 9 29 19 10 25 12 Sill in 25 02 12 15 27 t 34 29 14 40 15 15 24 6 12 Bt ou 28 21

Bt in 16 25 30 20 25 4 N 17 13 15 8 6 0 0 40 12 10 13 regional metamorphism ) 660000 18 2 454500 454000 5 15 5 A 12 regional metamorphism ) 30 25 2 40 2 0 60 12 12 18 453000 9 17 10 12 50 11 19 20 Contact (lithological) Fault (high-angle) Juzbado-Penalva do Castelo fault LGD-HGD boundary (Pinhel shear zone Biotite in (M Sillimanite in (M 60 12 15 13 14 14 30 23 14 40 35 30 4 25 15 38 45 15 20 22 7 15 10 25 28 45 25 7 11 22 10 11 11 40 21 15 7 35 22 21 22 10 12 26 12 18 14 6 20 19 15 8 20 i). Universal Transverse Mercator coordinates (zone 29). Princ. Dir.—principal direction. (E) Stereographic plot in the southern plot in the southern (E) Stereographic direction. Dir.—principal Princ. 29). (zone coordinates Mercator Transverse Universal i). 30 3 15 12 20 15 11 20 11 24 650000 13 10 ) 5 3 N 18 (295°/6°) (277°/3°) (300°/7°) NORT H 13 Location of sample for U-Pb geochronology TP- 4 m) 32 3 24 5 3 31 21 TP- 4 15 MEAN PRINC. DIR. MEAN PRINC. DIR. MEAN PRINC. DIR. 8 15 s 0 13 ) 28 INTERSECTION LINEATIO 11 0 451000 14 30 CRENULATION LINEATION (F MINERAL LINEATION (L -S upright folds converging to the southeast into the high-strain domain of the strike-slip shear zone. (D) Stereographic plot with the linear features plot with the linear features (D) Stereographic shear zone. domain of the strike-slip the high-strain the southeast into to converging folds upright 36 3 3 3 3 7 ] 30 (n=10) D D S (n=101 (n=4 ) 53 22 33 12 4 in the northern part of the Trancoso-Pinhel region (lower hemisphere, equal angle), including crenulation lineation (F including crenulation equal angle), hemisphere, (lower region Trancoso-Pinhel of the part in the northern 3 39 12 ° 15 e), and intersection lineation (L and intersection e), 28 3 16 25 30 36 29 15 33 25 30 15 26 30 contour lines High-strain domain of the Huebra shear zone Axial trace of D3 upright fold 26 0 16 15 640000 15 1% 2% 4% 8% 16% 32% 16 452000 ROSE (n=115 ) Sector angle = 10° Maximum = 50,4% [58 data Mean = 115°-295 25 52 D C B part of the Trancoso-Pinhel region. Trancoso-Pinhel of the part Figure 5. (A) Geological map of the northern part of the Trancoso-Pinhel region including the axial traces of folds and the main linear features (lithological legend from Fig. Fig. from (lithological legend and the main linear features of folds including the axial traces region Trancoso-Pinhel of the part (A) Geological map of the northern 5. Figure Trancoso- of the part (B) Geological map of the southern synf—synform. sync—syncline; Ant—antiform; also indicated. are geochronology U-Pb Locations of samples for 4). structure fan-like A) highlighting the zone (based on shear Huebra (C) Simplified map of the sinistral domain. HGD—high-grade domain; LGD—low-grade Pinhel region. of D of the axial traces associated with D associated tion lineation (L

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A B

C D

E F

Figure 6. (A) Tight D3 folds affecting granitic and pegmatitic dikes that intrude the metasedimentary rocks (to the left) along the northern boundary of

the Mêda-Escalhão massif (to the right). Ductile pervasive deformation in this area is attributed to the D3 Huebra shear zone. (B) Open D3 folds observed

in bedrock (S0, dashed yellowish lines) ~300 m away from the high-strain zone of the D3 Huebra shear zone. (C) Nearly isoclinal D3 folds affecting gra-

nitic and pegmatitic dikes (black dashed line) that intrude the metasedimentary rocks observed along the high-strain domain of the D3 Huebra shear zone. (D) Penetrative linear fabric of the metasedimentary rocks located along the northern boundary of the Mêda-Escalhão massif. The intense crenu-

lation lineation is associated with the development of numerous upright D3 microfolds (see intensively folded vein). (E) Crenulation cleavage (S4) and

crenulation lineation (F4) related to the Juzbado–Penalva do Castelo shear zone (JPSZ) in the northern limb of the Marofa synform. D4 folds affect S3. (F)

Mylonites of the high-strain domain of the D4 JPSZ. The presence of some feldspar porphyroclasts suggests that the shear zone is affecting a granitoid.

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To the north and south of the JPSZ, D3 structures recover their regional aligned quartz (ribbons) and muscovite. Grain size is significantly smaller

N130°–140°E trend. This is evident at a large scale (analysis of D3 fold for the dynamically recrystallized grains that, together with reoriented axial traces; Fig. 2) and can be also observed in the foliation map of the original igneous grains, define the main mylonitic foliation. Reoriented Trancoso-Pinhel region (Fig. 4B). The orientations of the structures in the grains, particularly muscovite, have larger size, intracrystalline defor- study area are controlled by the N80°–90°E to N60°–70°E trend of the mation (undulose extinction), and micro–kink folds. K-feldspar shows JPSZ. There are two contrasting structural imprints associated with the textures related to granular flow, such as the development and migration

JPSZ (D4); an early ductile deformation (D4.1), and a later brittle one (D4.2). of subgrains as well as partial transformation into plagioclase. K-feldspar We relate these two types of structures to the deformation induced by the may also show sigma shape and usually displays twins and symplectite.

JPSZ because they affect S3, they only occur adjacent to this shear zone, Plagioclase is fractured and displays arched twins. C shear planes may and their shear planes are systematically parallel to the trace of the JPSZ. be defined by chlorite, suggesting low-grade metamorphic conditions for

the development of S4.

Early Deformation Related to the JPSZ (D4.1) With the exception of the D4 shear bands, field observations indicate

One type of D4 ductile structure is a subvertical and spaced crenulation that this granitic massif is not affected by previous phases of deformation,

cleavage (S4) that is irregularly distributed in the tectonic block located to the and they have not been detected by petrographic analysis in thin section.

south of the JPSZ (Fig. 6E). This cleavage overprints S0–S3. The local short- This observation is significant in the study area, where D2 and/or D3 pro-

ening related to S4 is responsible for some of the folded pattern exhibited by duce a penetrative but variably intense planar fabric in the synorogenic

S3 along the core of the Marofa synform. The interference between this D3 granitoids (Díez Fernández and Pereira, 2016). These observations and

fold and larger folds associated with D4 crenulation produces a hook-type the elongate nature of this granitic massif parallel to the trace of the JPSZ

fold pattern at map scale (Fig. 4D). The apparent asymmetry of macro-D4 indicate that its intrusion can be considered synkinematic relative to D4.

folds is due to the south-dipping geometry of S3 in this area. D4 folds and

fabrics diminish in intensity to the south of the JPSZ as S3 becomes steeper Late Deformation Related to the JPSZ (D4.2)

and has a more oblique trend relative to the trace of the Marofa synform. D4 brittle deformation is represented by a narrow but poorly exposed

Another style of D4 ductile deformation is represented by a subvertical fault zone that cuts and removes portions of the main D4 mylonitic band.

mylonitic to ultramylonitic band (S4) that occurs in the quartzites defining The fault zone is at least 20 m wide, dips to the north, and consists of a the northern limb of the Marofa synform (Fig. 6F); the long subvertical complex set of gauges and breccias (Fig. 9A). The fault trace is parallel fault here may approximate the location of a mylonitic shear band (Fig. to the ductile JPSZ and to the mylonitic shear band that defines the core 4B). Although these two structures are not strictly the same, they are of the JPSZ; therefore, this fault is considered part of the same system. intimately related at a regional scale. The width and internal structure of However, deformation conditions were colder and the fabrics indicate this shear band in the study area are unknown because exposures are lim- that this is not a strike-slip fault. ited, of poor quality (Fig. 7A), and reworked by the fault (see following). A preliminary analysis of kinematic criteria in the fault zone did not

D4 shearing is progressively less intense and heterogeneous with provide clear indicators, in part because the fault zone cuts across a wide increasing distance from the mylonitic shear band located along the variety of rocks (phyllites, gneisses, migmatites, and granites) and tectonic

northern limb of the D3 Marofa synform, which probably represents the fabrics, with different pre-fault kinematic histories (D1–D4). Some dip-slip

core of the JPSZ. Other conjugate D4 shear bands broaden around the lineations and striations suggest a dominant dip-slip movement, although metasedimentary (migmatitic) host of the São Pedro–Vieiro granitic mas- in some sections shallowly plunging lineations (plunging 20°–50° to the sif (Fig. 7B), located in the northern limb of the Marofa synform. This west) may correspond to previous strike-slip deformation. Whether this is

granitic massif is also deformed by D4 shear bands (Fig. 7C), so the extent a normal or a reverse fault can be solved using a regional approach. The

of D4 ductile deformation to the north of the Marofa synform is ~2 km. lack of a wide LGD to the north of this fault, and the juxtaposition of the

In the metasedimentary rocks, S4 is a low-grade foliation defined by HGD and the Marofa synform cored by low-grade metasedimentary rocks

the reorientation of pre-S4 mineral grains into parallelism to the cleavage can be explained by an upthrown movement of the northern block in the or shear planes, and by a newly formed paragenesis consisting of fine- hanging wall of this fault, implying that it is a north-dipping high-angle

grained quartz + white mica + chlorite + opaque minerals. S4 in the shear thrust (Fig. 4D; Pereira et al., 2014). Coeval lateral movements or later bands displays sinistral kinematic criteria (Figs. 7B, 7C) compatible with brittle deformation and reworking cannot be ruled out.

the deflection of major structures at a regional scale. S4 follows the local Another type of late deformation is represented by a crenulation cleav- trend of the JPSZ, and is vertical or steeply dipping to the south (Fig. 8A). age that has a low to moderate dip and subhorizontal crenulation lineation.

S4 contains a shallowly plunging, N80°–90°E to N60°–70°E elongation This fabric produces open mesofolds that interfere with D3 folds (Fig. 9B). lineation defined by quartz and mineral aggregates, with plunge varying The only crosscutting relationship observed indicates that this is a post-

from the west to the east (L4e; Figs. 7D, 7E, and 8B). D4 crenulation (F4) D3 event, but no criteria are available to assess its timing relative to the

shows the same overall orientation and plunge variation as L4e. fabrics associated with the JPSZ or the later events (described in the following). The distribution of this type of deformation is random and local. São Pedro–Vieiro Massif: Petrography and Structure

The São Pedro–Vieiro massif is an elongated body of muscovite gran- Steeply Plunging Folds (D5) ite; the long axis in map view is parallel to the JPSZ (Fig. 4B). The

granitoid body consists of quartz, K-feldspar, plagioclase, muscovite, The shear bands related to D4 (Fig. 9C) and previous structures are

and minor tourmaline, apatite, chlorite, zircon, opaques, and some rare variably affected by north-south subvertical folds (D5; Figs. 7D and 9D).

biotite. Most of this massif is affected by sinistral D4 shear bands (Fig. These folds occur together with a nonpenetrative subvertical crenulation

7C), although there are sections of low strain and even undeformed rock cleavage (S5; Figs. 8C and 9E) and a crenulation lineation (F5; Figs. 7D where primary igneous textures are preserved. and 9D) that plunges between 30° and 80° to the south and southeast,

S4 in this granitoid body is formed in relation to sinistral C-S struc- largely controlled by the local orientation of S0–S4 (Fig. 8D). The ampli-

tures at solid-state conditions (Fig. 7F). C and S planes are defined by tude of mesoscale D5 folds may range between 2 and 3 m (Fig. 7D). These

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A B

C D

E F L4e

Figure 7. (A) Mylonites of the high-strain domain of the D4 Juzbado-Penalva do Castelo shear zone (JPSZ). These two examples are from quartzitic

layers of the northern limb of the Marofa synform. (B) Mylonites of the D4 JPSZ developed on paraderived migmatites. Note the sinistral kinematics

defined by extensional shear bands (C’) and sigma objects (felsic segregates derived from previous partial melting). (C) S-C shear planes in the syn-D4

São Pedro–Vieiro massif (muscovite granite) affected by the JPSZ. Note the sinistral kinematics. (D) D5 vertical folds affecting a low-grade foliation

associated with the D4 Juzbado-Penalva do Castelo shear zone. (E) Enlargement of part of D; note the development of a subhorizontal elongation

lineation (L4e) and a subvertical crenulation lineation (F5) related to the D4 JPSZ and the vertical folds, respectively. (F) D4 sinistral shear bands of the São Pedro–Vieiro granitic massif. Note the grain-size reduction of quartz, feldspar, and mica along the bands and the reorientation of igneous minerals (larger size) to foliation planes.

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ROSE (n=36) D4 CRENULATION LINEATION (F4) Sector angle = 10° (n=36) MEAN PRINC. DIR. (274°/2°) 3% contour lines Maximum = 30,6% [11 data] 6% 48% Mean = 65°-245° 12% 24% 3% contour lines A 6% B 12% 16%

D4 FOLIATION (S4) (n=36) POLES MEAN PRINC. DIR. (158°/84°) PLANES MEAN PRINC. DIR. (68°/84S°) ROSE (n=36) Sector angle = 10° Maximum = 36,1% [13 data] Mean = 68°-248° D4 ELONGATION LINEATION (L4e) (n=8) DX4 D5 3% contour lines 6% 12% 24% C

D5 CRENULATION LINEATION (F5) (n=45) MEAN PRINC. DIR. (167°/40°) D 3% contour lines 6% 12% 24%

ROSE (n=44) Sector angle = 10° Maximum = 22,2% [10 data] D5 FOLIATION (S5) (n=44) Mean = 170°-350° POLES MEAN PRINC. DIR. (80°/90°) PLANES MEAN PRINC. DIR. (170°/90°)

Figure 8. Stereographic plots. (A) The strike of S4. Princ. Dir.—principal direction. (B) The trend and variable plunge of D4 crenulation lineations

(F4). (C) The strike of S5. (D) The trend and plunge of D5 crenulation lineations (F5) (lower hemisphere, equal angle). Note that the dispersion of

S4 and F4 can be explained by rotation about D5 fold axes.

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A B

breccias

3 crenulation Post-D

breccias gauges C D

E F

Figure 9. (A) Breccias and gauges in the fault zone that cuts across the Juzbado-Penalva do Castelo shear zone. (B) Shallowly dipping, post-D3 crenulation

affecting D3 upright folds. (C) D4 sinistral shear bands located along the São Pedro–Vieiro massif (dashed black line) and affected by D5 moderate-plunging

folds. Note the development of a D5 crenulation lineation (F5). (D) Steeply plunging D5 folds developed in quartzites, which show penetrative crenula-

tion lineation (F5). (E) Thin section of a phyllite affected by D5 folds. The main fabric is a composite foliation (S0 + S1 + S3). The thin section is rotated 45°

clockwise from its original position. (F) Interference between D4 (east-west) and D5 (north-south) upright folds in the Ciudad Rodrigo area (after Díez

Fernández et al., 2013). Note the two contrasting crenulation lineations. These folds are nucleated from a very shallowly dipping S2 foliation, which

confers a shallowly plunging character to the D5 folds.

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folds occur together with centimeter-scale folds (Fig. 9D), which may be mass 207. The primary beam, composed of 16O16O+, is set to an intensity accompanied by a close-spaced crenulation (Fig. 9E). of ~5 nA, with a 120 μm Kohler aperture, which generates 17 × 20 µm

D5 folds occur throughout the Trancoso-Pinhel region, and are par- elliptical spots on the target. The secondary beam exit slit is fixed at 80 ticularly concentrated in places where the lithological units and the local μm, achieving a resolution of ~5000 at 1% peak height. structural elements define very open map-scale vertical folds. For example, All calibration procedures are performed on the standards included the trace of the JPSZ has a wavy pattern in the Trancoso-Pinhel region. on the same mount. Mass calibration is done on the REG zircon (ca. 2.5 The trace varies from N70°–80°E to N100°–110°E and then to N60°– Ga; very high U, Th, and common Pb content). Every analytical session

70°E (Fig. 4B). D5 folds are more abundant around areas showing these started measuring the SL13 zircon, which is used as a concentration kind of north-south inflections and therefore we consider them as major standard (238 ppm U). The TEMORA-2 zircon (416.8 ± 1.1 Ma, Black

D5 folds and responsible for much of the undulate structural trend of the et al., 2003), used as an isotope ratios standard, was then measured every region. Moreover, the length of the limbs that define these inflections is four unknowns. Data reduction and age calculations were done with the different at a regional scale. The two main inflections depicted by the SHRIMPTOOLS software (www.ugr.es/fbea). Crystallization ages and Marofa synform and the JPSZ define an open, Z-shaped vertical fold. Such concordia plots were obtained using Isoplot (Ludwig, 2008). asymmetry is compatible with a north-south dextral shearing. Results Late Faults Results of U-Pb SHRIMP analyses are listed in Table DR1 and pre- The latest phase of deformation is a set of subvertical faults that cuts sented in Figure 10. The uncertainties for individual analyses in the text, all the previous structures and rocks, including the latest Variscan gran- Table DR1, and concordia diagrams are given at the 1σ level, whereas itoids (Figs. 4A, 4B). Even the largest faults show limited offsets of the the uncertainties on weighted mean 206Pb/238U ages in the text and Figure regional tectonometamorphic structure. The fault set is divided into two 10 are given at the 2σ level (uncertainties have 95% confidence limits). principle families of approximately north-northeast–south-southwest and Cathodoluminescence images of the majority of zircons show complex northwest-southeast trends. The group trending north-northeast–south- internal structures. Composite grains have cores (xenocrysts) with oscil- southwest dominates in the northern part of the study area. latory and banded zoning, accompanied by dark or light rim overgrowths of variable widths (20–70 µm). Most of the analyses were performed on SHRIMP U-Pb ZIRCON DATING dark overgrowths. Few simple grains show oscillatory zoning.

The Almeida-Malpartida and São Pedro–Vieiro massifs were selected Foliated Granite for a U-Pb geochronological study. These two massifs are good candidates We obtained 18 U-Pb SHRIMP analyses for sample TP-4. Ages were for obtaining an absolute age reference for the strike-slip shear zones of the calculated assuming 204Pb correction. The older two zircon grains have Cryo- region. Dating of the nonfoliated Almeida-Malpartida biotite-rich granite genian (693 Ma) and Tonian (961 Ma) ages, suggesting inheritance. The constrains the age of the third deformation phase. This massif is discordant remaining 16 spots in single grains and in rim overgrowths of composite

across the Pinhel shear zone and the D3 axial traces recognized in the host grains spread along the concordia curve from ca. 352 to 287 Ma, and give a 208 238 metamorphic rocks (Fig. 4B). Dating of the foliated and syn-D4 São Pedro– weighted mean Pb/ Th age of 302.2 ± 8.3 Ma (mean square of weighted Vieiro muscovite-rich granite provides an age for the movement along the deviates, MSWD = 2.4; Fig. 10A). Some of the spread observed could be due

JPSZ, which, as discussed here, was developed after D3 and before D5. to the presence of antecrysts and/or caused by the combination of inheritance and Pb loss. In this group of grains, 7 grains in the age range of ca. 332–296 Analytical Technique Ma gave a weighted mean 208Pb/238Th age of 309 ± 11 Ma (MSWD = 0.6; Fig. 7A). A concordia age of 307.8 ± 3.1 Ma (MSWD = 1.8; Fig. 10A) is Two samples of granite (TP-4, São Pedro–Vieiro; TP-13, Almeida- considered the best estimate for the crystallization age of the foliated granite. Malpartida; see location in Fig. 5) were crushed, disk-milled, sieved (300 μm), concentrated on a Wilfley table, and separated via heavy liquids Nonfoliated Granite (bromoform and methylene iodide) at Universidade de Évora (Portugal) We performed 25 U-Pb SHRIMP analyses on sample TP-13. Ages were and Universidad Complutense de Madrid (Spain). Care was taken to select calculated using uncorrected ages; 6 of the zircon grains yield 206Pb/238U representative grains from fractions using no magnetic discrimination. ages of ca. 569–491 Ma that probably represent xenocrysts. A weighted Zircon grains were hand-picked under binocular microscope; the most mean 206Pb/238U age of 296.0 ± 7.5 Ma with a very poor fit (MSWD = 49; transparent were selected to calculate protolith crystallization ages. At Fig. 10B) is obtained with the remaining 20 spots, quite scattered in the the IBERSIMS Lab (SHRIMP Ion-Microprobe Laboratory of the Uni- concordia curve from ca. 336 to 281 Ma. Most of the scattering corresponds versity of Granada), zircon grains of each sample plus several grains of to analyses performed on zircon overgrowths that probably formed dur- standards were cast on a 3.5-cm-diameter epoxy mount, polished, and ing hydrothermal activity related to late magmatic processes. A group of documented using optical (reflected and transmitted light) and scanning 6 analyses in the age range of ca. 313–300 Ma yields a weighted mean electron microscopy (secondary electrons and cathodoluminescence). 206Pb/238U age of 304.0 ± 3.0 Ma (MSWD = 0.57; Fig. 10B). That age coin- After extensive cleaning and drying, mounts were coated with ultrapure cides with the age obtained from the upper intercept of an apparent common gold (8–10 nm thick) and inserted into a SHRIMP IIe/mc for analysis. Pb discordia line on concordia at 303.0 ± 2.1 Ma (MSWD = 0.89; Fig. 10B),

Each selected spot was rastered with the primary beam for 120 s prior to which is taken as the probable crystallization age of the post-D3 granite. the analysis, and then 6 scans were analyzed, following the isotope peak 196 204 204.1 206 207 208 238 248 sequence Zr2O, Pb, background, Pb, Pb, Pb, U, ThO, DISCUSSION 254UO. Every mass in every scan is measured sequentially 10 times with the following total counting times per scan: 2 s for mass 196; 5 s for The analysis of orientation, kinematics, and timing of strike-slip shear masses 238, 248, and 254; 15 s for masses 204, 206, and 208; and 20 s for zones is presented here as a tool for constraining finite paleostrain and

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data-point error ellipses are 68.3% conf 0,22 206Pb/238U 0,18 A 1000 206Pb/238U Concordia Age = 307.8 ±3.1 Ma (1σ, decay const. errors included) 0,060 0,14 MSWD (of concordance) = 1.8 800 Probability(of concordance) = 0.18 0,056 0,10 600 Sample TP-4 340 (foliated granite) 0,052 300 0,06 400 0,048 300 207Pb/235U 0,02 0,044 0,20,6 1,01,4 1,82,2

260 data-point error symbols are 2σ 0,040 data-point error ellipses are 68.3% conf 207 235 Age (Ma) Pb/ U Mean = 309±11Ma ( 95% conf.) 0,036 Wtd by data-pt errs only, 0 of 7 rej. 0,26 0,30 0,34 0,38 0,42 0,46 100 µm MSWD = 0.60, probability = 0.73 360

320 303 Ma 316 Ma 280

240 Figure 10. (A) U-Pb results of the Mean = 302.2±8.3Ma (95% conf.) syn-D4 São Pedro–Vieiro muscovite- Wtd by data-pt errs only, 0 of 16 rej. rich granite (TP-4). MSWD—mean MSWD = 2.4, probability = 0,002 square of weighted deviates; conf.—confidence; Wtd—weighted; pt.—point; rej.—rejected. (B) U-Pb data-point error ellipses are 68.3% conf results of the postkinematic 307 Ma 206 238 Almeida-Malpartida biotite-rich 0,095 Pb/ U 560 granite (TP-13). U-Pb concordia and 300 Ma mean age plots show all analyses 0,085 B 313 Ma performed that are younger than 480 ca. 352 Ma. 0,075

0,065 400 Sample TP-13 360 0,055 (non-foliated 100 µm 320 granite)

0,045 280 207Pb/235U 0,035 0,25 0,35 0,45 0,55 0,65 0,75 0,85

206Pb/238U data-point error symbols are 2σ 0,052 Age (Ma) Mean = 304±3,0Ma (95% conf.) 340 320 Interceptsat Wtd by data-pt errs only, 0 of 6 rej. 303,0±2.1Ma MSWD = 0.57, probability = 0.72 0,050 MSWD = 0,89 310 320

0,048 300 300

0,046 290 data-point error ellipses are 68.3% conf 280 207Pb/235U 0,044 0,10,3 0,50,7 0,91,1 1,3 Mean = 296.0±7.5Ma (95% conf.) Wtd by data-pt errs only, 0 of 20 rej. MSWD = 49, probability = 0,000

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inferring relative plate movements in orogeny. A distinction between structures sharing geometrical features and kinematics along with a cor- A relation of structures with contrasting geometry and kinematics (upright high-drag folds and strike-slip shear zones) have been made based on field criteria zone (crosscutting relationships), isotopic dating, and regional analysis.

Relative Timing of Strike-Slip Shear Zones

S4 S5 The Huebra and Tamames shear zones coincide with the formation of D3 folds. The Huebra shear zone has regional trend and kinematics similar to those of the JPSZ. Consequently, both strike-slip shear zones could have been formed together. However, our study shows that the structures B associated with D (at least the upright folds) are deflected and locally S 3 3 folded by the JPSZ (D ). In central Iberia, such relative timing between upright fold 4 strike-slip structures was previously described in the eastern section of N the JPSZ (e.g., Villar Alonso et al., 1992), where even the interference

between north-south–trending D3 folds and later east-west–trending D4 s folds associated with the JPSZ can be observed (Fig. 9F; Díez Fernández et al., 2013). We cannot exclude the possibility that the Huebra shear zone Huebra Malpica-Lamegoshear and zone remained active or was reactivated during subsequent sinistral deforma- Tamames shear zone tion associated with the JPSZ, because their shear planes almost coincide at a regional scale. In addition, minor sinistral strike-slip shear zones Continental convergence trending N70°–120°E have been recognized to the east and southeast of the Trancoso-Pinhel region, and are considered as post-D2 and postdating development of the Tamames synform (Díez Balda, 1986). The Tama- s mes synform has been interpreted as a D3 structure (Díez Fernández and C Pereira, 2016), so we propose that the aforementioned sinistral shear zones

represent yet another example of D4 strike-slip structures. Juzbado-Penalva do The JPSZ can be mapped northwest into the so-called Douro-Beira Castelo shear zone shear zone (Fig. 2). This alignment defines a single D strike-slip shear S4 4 zone that is deflected clockwise, from a northeast-southwest trend in the eastern part to a northwest-southeast trend in the western part. S upright folds 3 The fifth phase of deformation (D5) in the Trancoso-Pinhel region pro- N duced a series of north-south–trending folds accompanied by subvertical crenulation cleavage and steeply plunging fold axes. Equivalent folds are Continental convergence irregularly distributed throughout basement areas nearby and, in some cases, are accompanied by minor north-south– to north-northeast–south- D southwest–trending dextral strike-slip shear zones (e.g., Jiménez Ontive- ros and Hernández Enrile, 1983; Gil Toja et al., 1985; Díez Balda et al., 1990b; Díez Montes and Gallastegui, 1992; Villar Alonso et al., 1992; clockwise deflection Valle Aguado et al., 2000; Mellado et al., 2006). D5 folds affect the trace high-plunging

folds of JPSZ, and formed by east-west shortening. Previous studies consid- t ered these folds to be responsible for accommodating lateral movements S4 during the last pulses of the JPSZ (e.g., Valle Aguado et al., 2000). That

N interpretation is inconsistent with the fact that D5 folds actually affect the study strike-slip shear zone with which they are supposed to be related (Figs. 7D area S3 and 9C). Therefore, we provide an alternative interpretation for these folds. Continental convergenc Kinematic models of partitioned transpression postulate that the simple S5 shear and pure shear components of strain are decoupled into zones of

e faul Porto-Tomar deformation (Dewey et al., 1998). This way the simple shear component is preferentially accommodated by strike-slip faults, whereas the pure shear component produces shortening within the intervening tectonic Figure 11. (A) Model of partitioned transpression (after Teyssier et al., 1995) blocks (Fig. 11A; Teyssier et al., 1995; Dewey et al., 1998). Rotation of adapted to the Porto-Tomar shear zone. (B–D) Simplified kinematic models intervening blocks in response to the displacement along their bounding for the western part of the Central Iberian Zone showing the relative tim- strike-slip faults is expected (Mount and Suppe, 1987). The Porto-Tomar ing between late Carboniferous strike-slip shear zones. The development shear zone is a north-south–trending strike-slip shear zone associated with of the shear zones occurred under dextral partitioned transpression. These lateral dextral movement and east-west subhorizontal shortening, the latter models do not require important changes in the orientation of continental convergence, and may take place during protracted north-northeast–south- manifested as reverse faults and local upright folds (Ribeiro, 1974; Ribeiro southwest convergence (present-day coordinates). See text for discussion. et al., 1980). This shear zone drags and cuts D3 and D4 structures (Fig. 2).

(B) D3 Huebra, (C) D4 Juzbado-Penalva do Castelo. (D) D5 Porto-Tomar. At a large scale, the Porto-Tomar shear zone is considered responsible

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for the clockwise deflection of the Juzbado-Penalva-Douro-Beira shear and U-Pb isotopic dating of synorogenic granitoids (Díez Fernández and

zone (Martínez Catalán, 2011b), and deflection of other strike-slip shear Pereira, 2016). The age of D3 broadens a little if the age obtained for strike-

zones of southwest Iberia, such as the Coimbra-Córdoba shear zone (Fig. slip deformation along the D3 Malpica-Lamego shear zone (ca. 310–307 1; Pereira et al., 2010). Ma; Rodríguez et al., 2003; Gutiérrez-Alonso et al., 2015) and the age of

According to the relative timing, geometry, and kinematics of D5 struc- its synkinematic granitoids (ca. 320–314 Ma; Simões, 2000; Rodríguez tures in the study area (and in central Iberia in general), we propose that et al., 2007) are considered. All these ages are in agreement with the first

D5 is the far-field response to a broad east-west shortening related to absolute age estimations made for the tectonic fabrics associated with late

the Porto-Tomar shear zone. In our view, the Porto-Tomar shear zone is D3 upright folds of the Iberian Massif, dated as 314 ± 6 Ma (Rb-Sr and a partitioned transpressional shear zone, in which the high-strain zone K-Ar methods; Capdevila and Vialette, 1970; corrected by Ries, 1979). accommodated north-south dextral movements (simple shear), while The U-Pb age of ca. 303 Ma obtained for the postkinematic Almeida-

the reverse faults and D5 folds account for coeval east-west orthogonal Malpartida granite places an upper limit on the timing of D3 folding. The shortening (pure shear) distributed over the tectonic blocks bounded by age of 307.8 ± 3.1 Ma obtained for the São Pedro–Vieiro granite is con-

this shear zone (Fig. 11A). The clockwise deflection of former D3 and D4 sistent with the previously published K-Ar age of 305 ± 8 Ma (muscovite; structures along the Porto-Tomar shear zone resembles the model of high- Regêncio Macedo, 1988) for the same granitic body, and helps to constrain drag distributed shear proposed by Mount and Suppe (1987), so dextral the timing of the fourth deformation phase related to the movement in strike-slip deformation was probably coupled into a broad shear zone for the JPSZ. These geochronological data are in agreement with an Ar/Ar

a time. An equivalent rotation of tectonic blocks can be observed for D3 date of 309 ± 2 Ma obtained from mica that defines the S4 foliation in the

and D4 strike-slip shear zones. D3 folds rotate toward the shear planes eastern section of the JPSZ (Gutiérrez-Alonso et al., 2015). The Aguiar

of the D3 Huebra shear zone (Fig. 5C), whereas those folds and related da Beira granite, dated as 304 ± 8 Ma (U-Pb in zircon; Costa, 2011), cuts

fabrics are also deflected by the D4 JPSZ (Figs. 2 and 4). Rotation associ- the JPSZ and all the deformed granitoids in the Trancoso-Pinhel region

ated with D3 and D4 is restricted to a relatively narrow section along the (Fig. 4B). A similar age of 304 ± 2 Ma was estimated for the crystalliza- shear zones, thus making it possible to observe the primary orientation tion of the Villavieja de Yeltes granite (Fig. 2) (Gutiérrez-Alonso et al., of structures away from the shear zones. 2011), the emplacement of which was controlled by the development of an Considering the discussion here, at least three pulses of strike-slip extensional structure related to a north-south shear zone formed after the movement along steeply dipping shear zones can be identified: (1) the JPSZ (Mellado et al., 2006). The Cota-Viseu granodioritic-monzogranitic Huebra and Tamames shear zones along with other shear bands of lesser massif (Fig. 2), which was dated as 305 ± 4 Ma (Valle Aguado et al., 2005),

regional extent (D3; Fig. 11B), (2) the JPSZ and its continuation into the also postdates the JPSZ. Assuming some minor overlap in the isotopic

Douro-Beira shear zone (D4; Fig. 11C), and (3) the Porto-Tomar shear ages, D4 can be restricted to a range ca. 311–304 Ma. However, given

zone (D5; Fig. 11D). Each of these shear zones is accompanied by its own the consistency and concordance of the age data, this interval could be set of upright folds and minor shear bands. further restricted to ca. 309–305 Ma. In shear zone networks the faults are not all simultaneously active. The Porto-Tomar shear zone has been considered a transcurrent fault Shearing can consist of a pulsating process associated with strong parti- that might have been active during a significant period of the Variscan tioning (e.g., Carreras, 2001; Carreras et al., 2010). Progressive shearing orogeny (Ribeiro et al., 1990; Dias and Ribeiro, 1993; Shelley and Boss- gives rise to interference patterns that do not imply polyphase tectonics ière, 2000; Chaminé et al., 2003). However, the fault trace and structural (e.g., intramylonitic folding). Thus, differences in radiometric ages or imprint associated with this strike-slip system are late orogenic (Ribeiro deduced relative timing based on structural analysis do not exclude the et al., 1980, 2007, 2016; Pereira and Silva, 2001; Martínez Catalán et al., possibility of partitioned deformation in space and time. Detailed struc- 2007; Pereira et al., 2010), and the sedimentary control exerted by this tural analysis carried out in the study region shows the existence of pro- fault over adjacent basins started in the Late Pennsylvanian (e.g., Gama gressive deformation histories where deformation phases overlap. Spatial Pereira et al., 2008; Machado et al., 2011). strain partitioning is suggested by the development of individual strike-slip The Buçaco basin (Fig. 2) is a pull-apart basin, the development and shear zones. However, deformation was also partitioned through time. filling of which are associated with the Porto-Tomar shear zone (Domin- Radiometric dating indicates differences in age that cover an interval of gos et al., 1983; Wagner, 2004; Flores et al., 2010). The basin is early time of ~15 m.y. (see following), and crosscutting relationships between Gzhelian (palynology data; Machado, 2010), i.e., ca. 304 Ma (Cohen strike-slip shear zones give the impression of a sequence of individual et al., 2013). Because the entire basin is folded and affected by normal

phases of deformation in the region, i.e., D3–D5. Consequently, the strike- faults and thrusts formed during movement of the strike-slip system, the slip shear zones of the Iberian Massif are not coeval, and likely formed Porto-Tomar shear zone must have been active after ca. 304 Ma. Accord- during progressive strike-slip deformation. ing to the relative timing of phases of deformation and the correlation

established in this work between D5 and the Porto-Tomar shear zone, the Absolute Timing of Strike-Slip Shear Zones age of the latter should be younger than ca. 309–305 Ma (proposed age

for D4). This age estimate is in agreement with the onset of the Buçaco

Any model to explain the postextensional (post-D2) evolution of the basin ca. 304 Ma, with the Ar/Ar age obtained from syntectonic mica Central Iberian Zone must take into account that ensuing late Carboniferous grown in foliated granite of the Porto-Tomar shear zone (307 ± 5 Ma; strike-slip deformation was stepped in space and time (as inferred from the Gutiérrez-Alonso et al., 2015), and with the overprinting of the strike-slip

sequence of post-D2 events presented in this study; D3, D4, D5). This strike- deformation on ca. 308 Ma granites (Pereira et al., 2010) from the south- slip deformation probably occurred in a short period of time (ca. 317–304 ern branch of the Porto-Tomar shear zone. Moreover, the development

Ma; Regêncio Macedo, 1988; Rodríguez et al., 2003; Valle Aguado et al., of north-south D5 strike-slip shear zones locally controls the emplace- 2005; Gutiérrez-Alonso et al., 2015; Díez Fernández and Pereira, 2016). ment of granitic magmas (e.g., Villavieja de Yeltes granite; Mellado et

Timing of D3 deformation has been constrained in the Trancoso-Pinhel al., 2006). The crystallization age of this granite (Gutiérrez-Alonso et al.,

region to between 317 ± 9 Ma (age of the youngest syn-D2–D3 granitoid) 2011) provides further evidence for the ca. 304 Ma age for the onset of

and 311 ± 6 Ma (age of a syn-D3 granitoid) by means of structural data the Porto-Tomar shear zone.

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Some (e.g., Martínez Catalán, 2011a) consider that the Porto-Tomar zones and folds (e.g., Sanderson and Marchini, 1984), the orientation of

shear zone extends to the crustal-scale transcurrent shear zones of the the convergence vector for D3 would approach a north-northeast–south- Armorican Massif, thus depicting the Ibero-Armorican arc (e.g., Ribeiro southwest direction (Fig. 11B). Similar conclusions can be reached for

et al., 2007). The age of this orocline has been constrained by means of D4 and D5 (Figs. 11C, 11D), so convergence did not change its orientation

paleomagnetic and U-Pb data to ca. 310–295 Ma (Weil et al., 2010; Gutiér- significantly from D3 through D5, but produced a sequence of structures rez-Alonso et al., 2011). The postkinematic Lavadores granite, located with different orientations that have contrasting kinematics. Deformation at the northern branch of the Porto-Tomar shear zone (Fig. 2), yielded a partitioning associated with convergence could have been focused along Permian age of 294 ± 3 Ma (U-Pb zircon; Martins et al., 2014), so the local rheological anisotropies, thus resulting in strike-slip shear zones age of the Porto-Tomar shear zone must be older. with different orientations. The intrusion of granitic batholiths, the existence of large-scale shear zones (either extensional or strike slip), and the Kinematic Analysis contrasting rheological behavior between thermally different zones are

the most likely factors controlling the orientation and location of D3–D5 The sequence of individual phases of deformation that emerge from our shear zones. The thermal reequilibration of the orogen was ongoing long

analysis has been used to derive the evolution of paleostrain and consider after D2, as indicated by the presence of syn-D3, and even syn-D4 and

the kinematics of the waning stages of the Gondwana-Laurussia collision. syn-D5 granitoids. Deformation registered during D3–D5 in the study area The relative timing between strike-slip structures observed in the Trancoso- took place under low-grade greenschist facies conditions (chlorite zone). Pinhel region favors kinematic models where structures formed progres- The presence of hotter and less viscous material underneath, capable of sively (e.g., Martínez Catalán, 2011a) and questions those that assume sourcing synkinematic granitoids, implies a rheologically inhomogeneous they all were formed at the same time (e.g., Gutiérrez-Alonso et al., 2015). crust. The trace of late strike-slip shear zones could be controlled by such The strike-slip shear zones studied in this paper are among the larg- heterogeneities. As an example of orogenic reworking and conditioned

est of the central part of the Iberian Massif, and their noncoeval nature deformation in the Trancoso-Pinhel region, the D3 Huebra shear zone

implies that these shear zones cannot be integrated directly in an extru- and D4 JPSZ are parallel to a D2 extensional shear zone occurring at both sion model featuring conjugated faults. The existence of conjugated faults, sides of a granitoid massif located in the core of a large-scale antiform i.e., major faults accompanied by minor subsidiary faults with oblique (i.e., Pinhel shear zone; Díez Fernández and Pereira, 2016). Therefore, trace and opposite kinematics, is possible for some shear zones of the the oblique trace of sinistral shear zones in the Iberian Massif relative to Iberian Massif (e.g., Iglesias Ponce de Leon and Choukroune, 1980; Díez major dextral shear zones is not necessarily imposed by the orientation Fernández and Martínez Catalán, 2012). Such a model was used as part of the stress field, as it could be expected for a conjugated fault system. of the working hypothesis in previous studies that attempted integration We suggest a control on the trace of major strike-slip shear zones exerted of all the strike-slip shear zones in a single tectonic process (e.g., Shelley by previous structures, which would have conferred a conjugated appear- and Bossière, 2002; Gutiérrez-Alonso et al., 2015). Those approaches ance to some of the late Variscan structures. need to be revised for the case of major shear zones such as Porto-Tomar, In the Late Devonian, convergence affecting the margin of Gond- Coimbra-Córdoba, Juzbado-Penalva do Castelo, and Malpica-Lamego, wana included a dextral component (oblique continental subduction; Díez among others, because these shear zones are not coeval. Fernández et al., 2012a). Lateral translations continued to be significant

The D3 Tamames and Huebra shear zones represent the onset of intrac- into the early Carboniferous, when the oblique convergence was coeval ontinental deformation after a period of orogenic extensional collapse with (1) the exhumation of high-pressure rocks, (2) the emplacement of

(D2; Díez Fernández and Pereira, 2016). The orthogonal components of allochthonous nappes (Díez Fernández and Martínez Catalán, 2012), (3)

ongoing convergence were distributed across the orogen and produced D3 the generation of a subduction-related tectonic mélange (Arenas et al., upright folds. In the northwest and central Iberian section of the orogen, 2009), and (4) the subsequent extensional collapse of the orogenic hinter- dextral strike-slip structures at this stage are more dominant than sinistral land (Díez Fernández et al., 2012b). The development of late strike-slip ones. The Tamames shear zone, and more important, the Malpica-Lamego shear zones under dextral transpression during the late Carboniferous shear zone (Figs. 1 and 2), which are several hundreds of kilometers long indicates the persistent character of right-lateral displacements between (Coke and Ribeiro, 2000; Llana-Fúnez and Marcos, 2001), are related to Gondwana and Laurussia throughout the late Paleozoic.

the development of D3 upright folds (Díez Fernández and Martínez Cata- lán, 2012; Pamplona et al., 2015) and in some instances were exploited Strike-Slip Shear Zones and the Iberian Oroclines by the intrusion of deep-sourced magmas (e.g., Gallastegui, 1993), thus indicating their crustal-scale nature. Sinistral shear zones located nearby The Ibero-Armorican arc is a late orogenic or postorogenic orocline are interpreted to represent subsidiary shear zones equivalent to bookself- that bends the entire Variscan orogen of Iberia (Weil et al., 2013), thus type structures (Iglesias Ponce de Leon and Choukroune, 1980; Shelley conferring a dominant northwest-southeast trend to the major structures and Bossière, 2000). Accordingly, and given that the trace of these shear that occupy its southern branch (Fig. 1). In that branch, the Central Iberian zones is parallel to the structural trend of the orogen, convergence likely arc represents another orocline that seems to be coupled to the Cantabrian included a right-lateral component, which was concentrated in strike- orocline (Aerden, 2004), the core of which is located to the north and slip shear zones such as Malpica-Lamego and Tamames, among others. convex in the opposite direction, i.e., representing a roughly equivalent The coexistence of late Carboniferous strike-slip shear zones and orogenic curvature that encompasses all of Iberia. Although the mechan-

widespread upright folds during D3–D5 suggests that they formed during ics, kinematics, and even the existence and geometry of these plate-scale partitioned transpression (Fig. 11A). The northwest-southeast orienta- vertical folds are still discussed (e.g., Martínez Catalán, 2011a; Shaw et

tions of D3 folds provide a constraint for the maximum shortening direc- al., 2012; Weil et al., 2013; Martínez Catalán et al., 2014; Pastor-Galán tion at this stage, i.e., northeast-southwest (in present-day coordinates) et al., 2015; Dias et al., 2016), the development of strike-slip shear zones assuming limited rotation of the folds during transpression. If the rela- is considered to be intimately related to the evolution of the Iberian oro- tive movement between intervening landmasses included a dextral lat- clines (e.g., Martínez Catalán, 2011a; Weil et al., 2013; Gutiérrez-Alonso eral component, and convergence was partitioned into strike-slip shear et al., 2015).

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Whereas the Ibero-Armorican arc affects many of the strike-slip shear REFERENCES CITED zones (note their curved pattern in Fig. 1), none of them conform to the Aerden, D.G.A.M., 2004, Correlating deformation in Variscan NW-Iberia using porphyroblasts; implications for the Ibero-Armorican Arc: Journal of Structural Geology, v. 26, p. 177–196, arched geometry of the Central Iberian arc. The JPSZ and Porto-Tomar doi:10.1016/S0191-8141(03)00070-1. shear zone, for example, cut across the axial zone of the Central Iberian Arenas, R., Sánchez Martínez, S., Castineiras, P., Jeffries, T.E., Díez Fernández, R., and An- arc without being affected by it. Accordingly, (1) the Ibero-Armorican arc donaegui, P., 2009, The basal tectonic melange of the Cabo Ortegal Complex (NW Iberian Massif): A key unit in the suture of Pangea: Journal of Iberian Geology, v. 35, p. 85–125. is the latest orocline formed in this part of the orogen (Martínez Catalán, Arenas, R., Díez Fernández, R., Sánchez Martínez, S., Gerdes, A., Fernández-Suárez, J., and 2011a); (2) the Ibero-Armorican arc was nucleated after ca. 309–304 Ma, Albert, R., 2014, Two-stage collision: Exploring the birth of Pangea in the Variscan terranes: Gondwana Research, v. 25, p. 756–763, doi:10.1016/j.gr.2013.08.009. which is the putative age of the onset of the Porto-Tomar shear zone (this Arenas, R., et al., 2016, Allochthonous terranes involved in the Variscan suture of NW Iberia: age is also supported by abundant paleomagnetic and structural data; Weil A review of their origin and tectonothermal evolution: Earth-Science Reviews, v. 161, et al., 2010, 2013); and (3) the Central Iberian arc was nucleated before p. 140–178, doi:10.1016/j.earscirev.2016.08.010. 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However, the Trancoso-Pinhel region is located along Carreras, J., Druguet, E., Griera, A., and Soldevila, J., 2004, Strain and deformation history one of the branches of the Ibero-Armorican arc and, consequently, the in a syntectonic pluton. The case of the Roses granodiorite (Cap de Creus, Eastern Pyr- original orientation of the vector of convergence for D –D must have enees), in Alsop, G.I., and Holdsworth, R.E., eds., Flow processes in faults and shear 3 5 zones: Geological Society of London Special Publication 224, p. 307–319, doi:10.1144 been rotated to some extent during orocline development. In any case, the /GSL.SP.2004.224.01.19. 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Therefore, major dextral and sinistral Dias, R., and Ribeiro, A., 1998, Interaction between major sinistral wrench faults and coeval folds in a Variscan transpressive regime (NE Portugal): Comunicações dos Serviços shear zones cannot be integrated in an extrusion model featured by (coeval) Geológicos de Portugal, v. 85, p. 19–27. conjugated faults. Even though the trace and kinematics of the analyzed Dias, R., Coke, C., and Moreira, N., 2010, Deformação Varisca heterogénea no eixo Marão–Foz strike-slip shear zones are remarkably different, no major changes in the Côa (autóctone da Zona Centro Ibérica); implicações para a estrutura regional: e-Terra, v. 11, p. 1–4. orientation of continental convergence are required to explain the sequence Dias, R., Ribeiro, A., Coke, C., Pereira, E., Rodrigues, J., Castro, P., Moreira, N., and Rebelo, of late Carboniferous structures observed. Such differences, however, can J., 2013, Evolução estrutural dos sectores setentrionais do Autóctone da Zona Centro- be explained by a rheologically heterogeneous lithosphere by the end of Ibérica, in Dias, R., et al., eds., Geologia de Portugal, Volume I: Geologia Pré-mesozóica de Portugal: Lisbon, Portugal, Escolar Editora, p. 73–147. the Variscan orogeny. Mechanical anisotropies were abundant after the Dias, R., Ribeiro, A., Romão, J., Coke, C., and Moreira, N., 2016, Reviewing the arcuate struc- gravitational and thermal reequilibration of the orogen, which produced tures in the Iberian Variscides; constraints and genetical models: Tectonophysics, v. 681, p. 170–194, doi:10.1016/j.tecto.2016.04.011. widespread magmatism and extensional shear zones that constituted weak Díez Balda, M.A., 1986, El Complejo Esquisto-Grauváquico, las series Paleozoicas y la estruc- zones for absorbing subsequent lateral movements upon superimposed tura Hercínica al Sur de Salamanca: Salamanca, Spain, Universidad de Salamanca, 162 p. compression. The widespread nature of late upright folds across the Iberian Díez Balda, M.A., and Vegas, R., 1992, La estructura del Dominio de los pliegues verticales de la Zona Centro-Ibérica, in Gutiérrez-Marco, J.C., et al., eds., Paleozoico Inferior de Ibero- Massif indicates a tectonic origin for such compression, likely derived from América: Extremadura, Spain, Universidad de Extremadura, p. 523–534. further convergence between Gondwana and Laurussia. A kinematic analy- Díez Balda, M.A., García Casquero, J.L., Monteserín López, V., Nozal Martín, F., Pardo Alonso, sis of the late strike-slip shear zones suggests that they all may have resulted M.V., and Robles Casas, R., 1990a, Cizallamientos subverticales posteriores a la segunda fase de deformación Hercínica al sur de Salamanca (Zona Centro-Ibérica): Revista de la from dextral oblique convergence. This type of convergence also explains Sociedad Geológica de España, v. 3, p. 117–125. the previous tectonic evolution recorded in central and northern Iberia Díez Balda, M.A., Vegas, R., and González Lodeiro, F., 1990b, Central-Iberian Zone. autochtho- nous sequences: Structure, in Dallmeyer, R.D., and Martínez García, E., eds., Pre-Mesozoic related to folds, thrusts, and extensional faults, and points out the long- geology of Iberia: Berlin, Germany, Springer-Verlag, p. 172–188. lasting nature of lateral movements during the amalgamation of Pangea. Díez Balda, M.A., Martínez Catalán, J.R., and Ayarza, P., 1995, Syn-collisional extensional collapse parallel to the orogenic trend in a domain of steep tectonics: The Salamanca detach- ment zone (Central Iberian Zone, Spain): Journal of Structural Geology, v. 17, p. 163–182, ACKNOWLEDGMENTS doi:10.1016/0191-8141(94)E0042-W. We thank Stephen Johnston and Brendan Murphy for insightful revisions. Financial support Díez Fernández, R., and Arenas, R., 2015, The Late Devonian Variscan suture of the Iberian was provided by Fundação para a Ciência e Tecnologia (Portugal) through the research project Massif: A correlation of high-pressure belts in NW and SW Iberia: Tectonophysics, v. 654, GOLD (Granites, Orogenesis, Long-term strain/stress, and Deposition of ore metals) (PTDC/ p. 96–100, doi:10.1016/j.tecto.2015.05.001. GEO-GEO/2446/2012: COMPETE: FCOMP-01–0124-FEDER-029192). Díez Fernández appreciates Díez Fernández, R., and Martínez Catalán, J.R., 2012, Stretching lineations in high-pressure belts: financial support from Fundação para a Ciência e Tecnologia (Portugal) through its postdoctoral The fingerprint of subduction and subsequent events (Malpica-Tui complex, NW Iberia): program (SFRH/BPD/85209/2012). This is IBERSIMS (SHRIMP Ion-Microprobe Laboratory of the Journal of the Geological Society [London], v. 169, p. 531–543, doi:10 .1144 /0016 -76492011 -101. University of Granada) publication 41, and a contribution to IGCP (International Geoscience Díez Fernández, R., and Pereira, M.F., 2016, Extensional orogenic collapse captured by strike- Programme) Project 648 (Supercontinent Cycle and Global Geodynamics). slip tectonics: Constraints from structural geology and U-Pb geochronology of the Pinhel

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