Tectonothermal evolution in the core of an arcuate fold and thrust belt: the southeastern sector of the Cantabrian Zone (Varis

The tectonothermal evolution of an area located in the core of the Ibero-Armorican arc (Variscan belt) 7 has been determined by using the conodont color alteration index (CAI), Kübler index of illite (KI), the Árkai 8 index of chlorite (AI), and the analysis of clay minerals and rock cleavage. The area is part of the Cantabrian 9 Zone (CZ), which represents the foreland fold and thrust belt of the orogen. It has been thrust by several large 10 units of the CZ, what resulted in the generation of a large amount of synorogenic Carboniferous sediments. CAI, 11 KI and AI values show an irregular distribution of metamorphic grade, independent of stratigraphic position. 12 Two tectonothermal events have been distinguished in the area. The first one, poorly defined, is mainly located 13 in the northern part. It gave rise to very low-grade metamorphism in some areas and it was associated with a 14 deformation event that resulted in the emplacement of the last large thrust unit and development of upright folds 15 and associated cleavage (S1).The second tectonothermal event gave rise to low-grade metamorphism and 16 cleavage (S2) crosscutting earlier upright folds in the central, western and southern parts of the study area. The 17 event continued with the intrusion of small igneous rock bodies, which gave rise to contact metamorphism and 18 hydrothermal alteration. The second event was linked to an extensional episode due to a gravitational instability 19 at the end of the Variscan deformation. This tectonothermal evolution occurred during the Gzhelian-Sakmarian. 20 Subsequently, several hydrothermal episodes took place, in association with local development of crenulation 21 cleavage during the Alpine deformation. 22


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The Variscan belt defines an arc in the northwestern Iberian Peninsula (Ibero-Armorican Arc), whose core is 24 formed by the Cantabrian Zone (CZ), which represents the foreland fold and thrust belt of the orogen (Fig. 1).

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This zone consists of Palaeozoic rocks in which two tectonostratigraphic units have been distinguished (Julivert,

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1978; Marcos and Pulgar, 1982), whose limit is approximately located by the Devonian-Carboniferous 27 boundary. The preorogenic unit is formed of Cambrian to Devonian rocks consisting of alternating carbonate and 28 siliciclastic formations; they form a wedge that thins towards the foreland. The synorogenic unit is formed by 29 several clastic units, also thinning towards the foreland, which filled foredeep basins generated in the front of the 30 main thrust units of the CZ. In this zone, the Variscan deformation occurred during the upper Carboniferous and 31 gave rise to thin-skinned tectonics, with several large thrust units and associated folds (Fig. 1); the units were 32 emplaced in a sequence towards the foreland. The deformation occurred under shallow crustal conditions, so that 33 diagenetic conditions are dominant in the zone and absence of cleavage in the rocks is also dominant.

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Fernández and Heredia, 1987), which occurred during the late Moscovian. In the same episode that involved the 75 emplacement of the Picos de Europa thrust unit during the Kasimovian-Gzhelian (Merino-Tomé et al. 2009), N -76 S shortening occurred in the study area, involving the development of thrusts, high angle reverse faults and the 77 reactivation of older faults with a movement dominantly southward (Maas, 1974). In addition, upright folds with 78 E -W axial trace developed; among them, the Curavacas-Lechada syncline is remarkable for its notable 79 dimensions (Savage, 1967;Lobato, 1977;Rodríguez Fernández, 1994).

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Preparation of samples and methods for XRD analysis follow the methods described in Brime et al. (2003) 129 Reaction progress in illitic minerals (sensu Środoń 1984) has been widely used to assess the evolution of

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∆ o 2 to minimize variations caused by differences in recording conditions. For this study the KI was measured 135 using a laboratory procedure similar to that outlined by the IGCP 294 working group (Kisch, 1991). The 136 numerical KI value decreases with improving "crystallinity" and is expressed as small changes in the Bragg 137 angle ∆ o 2, using Cu K radiation.

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For the problems involved in the use of the CIS scale to assess metamorphic grade see Brime (1999) and

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The KI method does not allow temperature constraints to be placed on the upper and lower boundaries of 152 the anchizone and it is more likely to be a measure of reaction progress than of the thermodynamic equilibrium 153 achieved (Essene and Peacor,1995). However, this method, in combination with others such as fluid inclusions 154 or reflectance of carbonaceous material, indicates that the transition diagenesis-anchizone could be correlated 155 with a temperature of 230 ± 10 ºC, whereas the limit anchizone-epizone would be at 300 ± 10 o C (Müllis, 1979; 156 Frey et al., 1980;Frey, 1987

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For the metamorphic zonation from CAI data, we use the terminology described by García-López et al.

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Temperature ranges of the CAI values were obtained from the Arrhenius plot presented by Epstein et al.

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(1977) and Rejebian et al. (1987). The maximum possible heating time is the age of the rock. Nevertheless, it is

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In the context of the study area, we call tectonothermal event to a deformational event with cleavage

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Grade ranges from deep diagenetic to epizonal, but deep diagenetic and mainly low anchizonal metapelites 267 are predominant in most of the areas (Figs. 6 and 7). Expandability of the 10 Å peak is only lost at the high 268 anchizone to epizone boundary. Deep diagenetic areas can be found to the north (Liébana and Valdeón areas).

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The Riaño-Cervera area is mainly low anchizonal with a few samples being diagenetic or deep anchizonal. In the 270 Pisuerga area, the grade ranges from deep diagenetic to low anchizonal (Figs. 6, 7). Higher grade (epizonal) 271 samples may appear in any formation and they are more abundant in the western part of the Yuso-Carrión area,

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where the Peña Prieta granodiorite is located, and in the Devonian of the Valsurbio area (Fig.7). In both cases, it 273 is in those high grade samples where chloritoid is more abundant (Fig.6).

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The existence of some discrepancies between KI and AI may be caused by the presence of small amounts   as suggested in the Glarus Alps by Frey (1978), who considered Prl an indicator of anchizonal regional 301 conditions. In fact of the 53 samples in which Prl is present, Kln is found, and in very small amounts, in only 8 302 of them. However, the stability field of Prl is strongly influenced by water activity, and thus the formation 303 temperature could be notably lower (Thompson, 1970;Winkler, 1979;Hemley et al.,1980). Its presence in 304 diagenetic samples is not uncommon and could be due to the influence of magmatic fluids (Hosterman et al.,

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Presence of Cld in the anchizone has been discussed by Kisch (1983), who concluded that Cld cannot

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The CAI values are in general independent of the stratigraphic position of the samples (Figs. 8 and 9). The

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The lack of carbonate rocks prevents in some areas the construction of a complete map of CAI isograds;

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The apparently chaotic distribution of CAI isograds could be due to a heat from subsurface intrusions at 367 depth resulting in isotherms having complex geometry. This pattern may be related to a crustal thinning during

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In some locations, S 2 is gently folded with local development of crenulation cleavage. This may be a result 415 of the Alpine deformation, which is the only post-Variscan compressional deformation described in the area 12 (Gallastegui, 2000 and references therein), and involved a ductile deformation that required a moderate 417 temperature and gave rise to a dome shape in the Valsurbio unit (Marín, 1997).

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The

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The correlation among the different indicators that can be used to establish the metamorhic grade is