Triassic redbeds in the Malaguide Complex (Betic Cordillera — Spain): Petrography, geochemistry and geodynamic implications

Earth-Science Reviews 117 (2013) 1–28

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Triassic redbeds in the Malaguide Complex (Betic Cordillera — Spain): Petrography, geochemistry and geodynamic implications Francesco Perri a,⁎, Salvatore Critelli a, Agustín Martín-Algarra b, Manuel Martín-Martín c, Vincenzo Perrone d, Giovanni Mongelli e, Massimiliano Zattin f a

Dipartimento di Scienze della Terra, Università della Calabria, 87036 Arcavacata di Rende (CS), Italy Departamento de Estratigrafía y Paleontología, Universidad de Granada, Avenida Fuentenueva s/n, 18071 Granada, Spain Departamento de Ciencias de la Tierra y del Medio Ambiente, Universidad de Alicante, Ap. 99 03080 Alicante, Spain d Dipartimento di Scienze della Terra, della Vita e dell'Ambiente, Università “Carlo Bo” di Urbino, Campus Scientifico, località Crocicchia, 61029 Urbino, Italy e Dipartimento di Chimica, Università degli Studi della Basilicata, Via N. Sauro, 85100, Potenza, Italy f Dipartimento di Geoscienze, Università degli Studi di Padova, 35137, Padova, Italy b c

a r t i c l e

i n f o

Article history: Received 13 January 2012 Accepted 9 November 2012 Available online 22 November 2012 Keywords: Redbeds Malaguide Complex Composition Provenance Geodynamic implications

a b s t r a c t Sandstone petrography and mudstone mineralogy and geochemistry of Triassic mudstones and sandstones from continental redbeds of the Malaguide Complex (Betic Cordillera, southern Spain) provide useful information on provenance, palaeoclimate and geodynamics during the early stages of the Pangea break-up, and on their diagenetic evolution. The sandstones are quartzarenites to sub-litharenites, with minor lithic fragments and rare feldspars. The mudstone samples show a PAAS like elemental distribution. The samples likely record recycling processes from their metasedimentary basement rocks that significantly affected the weathering indices, and monitors cumulative effects, including a first cycle of weathering at the source rocks. Sandstone composition and chemical–mineralogical features of mudstones record a provenance derived from continental block and recycled orogen that were weathered under warm and episodically wet climate. Source areas were located towards the east of the present-day Malaguide outcrops, and were formed by fairly silicic rock types, made up mainly of Palaezoic metasedimentary rocks, similar to those of the Paleozoic underlying series, with subordinate contributions from magmatic–metamorphic sources, and a rare supply from mafic metavolcanic rocks. Claymineral distribution of mudstones is dominated by illite and illite/smectite mixed-layer that result from differences in provenance, weathering, and burial/temperature history. Illite crystallinity values, illitization of kaolinite, occurrence of typical authigenic minerals and apatite fission-track studies, coupled with a subsidence analysis of the whole Malaguide succession suggest burial depths of at least 4–6 km with temperatures of 140–160 °C, typical of the burial diagenetic stage, and confirm the Middle Miocene exhumation of the Betic Internal Domain tectonic stack topped by the Malaguide Complex. © 2012 Elsevier B.V. All rights reserved.

Contents 1. 2.

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Introduction . . . . . . . . . . . . . . . . . . . . . . . Geological setting . . . . . . . . . . . . . . . . . . . . . 2.1. The Malaguide Complex . . . . . . . . . . . . . . . 2.2. The Saladilla Formation . . . . . . . . . . . . . . . Stratigraphy and facies of the studied Triassic sections . . . . 3.1. Western sector (Malaga) . . . . . . . . . . . . . . 3.2. Eastern sector (Sierra Espuña) . . . . . . . . . . . . Sampling and analytical methods . . . . . . . . . . . . . . Sandstone petrology and detrital modes . . . . . . . . . . 5.1. Interstitial composition . . . . . . . . . . . . . . . Mineralogical and geochemical composition of mudstones . . 6.1. Mineralogical results according to stratigraphic sections 6.2. Major and trace element variations . . . . . . . . .

⁎ Corresponding author. Tel.: +39 0984 493550. E-mail address: [email protected] (F. Perri). 0012-8252/$ – see front matter © 2012 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.earscirev.2012.11.002

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7. 8. 9.

Fission-track study . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Subsidence analysis and compaction . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.1. Provenance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2. Source-area weathering, and climatic features . . . . . . . . . . . . . . . 9.3. Sorting and recycling . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4. Malaguide Triassic palaeogeography, and evolution of the relief in the source 9.5. Diagenesis, and burial/thermal history . . . . . . . . . . . . . . . . . . 10. Palaeogeographic summary and concluding remarks . . . . . . . . . . . . . . . Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Appendix A. Subsidence analysis and compaction procedures . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1. Introduction Redbeds are important indicators of the paleogeographic, paleotectonic and paleoclimatic conditions controlling continental basins, preferentially when they develop in relation to early stages of rifting, which will lead to the break-up of a supercontinent. The sedimentological, petrographical, mineralogical and geochemical features of redbeds provide useful information on the sedimentary, climatic, diagenetic and tectonic processes characterizing the geodynamic and palaeogeographic evolution. These processes are testified by well known redbed examples of Precambrian (Walker, 1967; Turner, 1982;

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Gutzmer et al., 2006), Palaeozoic (Hofmann, 1991; Ludwig, 1992; Schöner and Gaupp, 2005) and Mesozoic age (Cassinis et al., 1979; López-Gómez and Arche, 1993; Arche and López-Gómez, 1996; Perrone et al., 2006; De Vicente et al., 2009, and references therein). These are usually interpreted as alluvial deposits related to weathering and erosion under warm tropical climates with frequent arid episodes, and deposited within rift-related extensional basins. Middle–Upper Triassic redbeds, in particular, constitute important stratigraphic markers characterizing the base of the Alpine cycle in many internal units of the Western-Central Mediterranean Alpine Chains, from the Betic Cordillera to the Northern Apennines (Roep,

Fig. 1. Schematic geological map of the Betic Cordillera, with indication of the study areas.

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1972; Martín-Algarra et al., 1995; Iannace et al., 2005; Perri et al., 2005a, 2005b, 2008a, 2008b, 2011; Mongelli et al., 2006; Perrone et al., 2006; Aldinucci et al., 2008; Critelli et al., 2008; Zaghloul et al., 2010). All these papers pointed out the importance of Triassic redbeds for the reconstruction of the paleogeographic, and paleotectonic evolution of the early stages of the disintegration of Pangea (see also Arche and López-Gómez, 1996; Baudon et al., 2009; De Vicente et al., 2009; McKie and Williams, 2009, among many others). Perrone et al. (2006) pointed out that the Alpine Triassic redbeds cropping out in the Internal Domains of the Western Mediterranean Chains were palaeogeographically and palaeotectonically independent of the European–African Germano-Andalusian Triassic redbeds cropping out in the External Domains of the same chains and in their forelands. Moreover, the same authors also stated that: 1) such redbeds were deposited around small erosional landmasses surrounded by shallow marine carbonates with Alpine lithofacies; 2) they constituted the continental to coastal depositional belt of a well-defined Triassic palaeogeographical domain that laterally evolved from erosional to shallow marine areas; 3) they made part of a (Mesomediterranean) Triassic continental crust block interposed between the future IberiaEurope, Africa and Adria–Apulia plates; 4) since Mid-Jurassic time, this block evolved to a Mesomediterranean Microplate that was detached from the Iberia-Europe, Africa and Adria Plates, and that remained isolated from them by narrow oceanic basins in the westernmost areas of the Alpine Tethys (Guerrera et al., 1993; Perrone, 1996;

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Bonardi et al., 2001; Michard et al., 2002); and, finally 5) this continental crust block later functioned as the hinterland of the Western Mediterranean Alpine Belts. In this contribution we present a detailed study of the petrographical, mineralogical and geochemical features, coupled with a subsidence analysis, of the Triassic continental red sandstones and mudstones characterizing the Malaguide Complex (Saladilla Formation) of the Betic Cordillera (southern Spain). The study is aimed to a better understanding of their composition, provenance and paleoclimatic signatures in the context of the early Mesozoic Triassic rifting of Pangea, and of the significance of their diagenesis in relation to the later evolution of the westernmost Mediterranean Alpine regions. 2. Geological setting 2.1. The Malaguide Complex Three main tectono-palaeogeographic domains have been recognized in the Betic Cordillera: Internal, Flysch and External Domains (Fig. 1). The Internal Domain forms a tectonic stack of both basement and cover Alpine thrust nappes. These are known as Nevadofilabride, Alpujarride and Malaguide complexes, plus a complex of imbricated Frontal Units (Fig. 1). These complexes formed after orogenic deformation of both the Nevadofilabride oceanic realm and the Mesozoic Mesomediterranean Microplate and its surrounding margins. The

Fig. 2. Synthetic geological map of the Malaga area showing outcrops of the Malaguide Complex and location of sections.

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External and Flysch Domains are formed by thin-skinned thrust nappes made of Meso-Cenozoic sediments. Both realms individuated after the Pangea break-off, when the Mesomediterranean continental block was definitely detached from the Iberia-Europe and Africa-Adria continental crust in Mid-Jurassic time (Vera, 2004). The Malaguide Complex constitutes the highest tectonic element of the Betic Internal Domain. It thrusts widely over both the Alpujarride Complex and the Frontal Units (Fig. 1). Stratigraphically, it mainly consists of pre-Ordovician to Upper Carboniferous clastic (meta)-sediments strongly deformed during the Variscan Orogeny (Martín-Algarra et al., 2009a). This basement is unconformably covered by a Meso-Cenozoic sedimentary cover (Figs. 2 and 3) that starts with well-developed Triassic continental redbeds (Fig. 4). The Malaguide Triassic continental facies belt changed laterally to marine carbonate sediments with Alpine

lithofacies of the Alpujarride Complex and of the Frontal Units. The latter made part of a carbonate platform that constituted the western extension of the Austroalpine, South-Alpine and Apeninic shallow marine Triassic realms (Martín-Algarra, 1987; Martín-Rojas et al., 2009 and references therein). Depending on the area considered, the Malaguide Complex is constituted by one to four different alpine tectonic units (Martín-Algarra et al., 2009b). In western outcrops (Malaga Mountain area: Fig. 2) one single tectonic unit is usually considered. Other Malaguide tectonic units can be recognized northwards, along the Internal–External Zone boundary (Martín-Algarra, 1987; Martín-Algarra et al., 2009b). The Malaguide Triassic redbeds, however, do not show any particular change in facies or thickness depending on the tectonic units considered.

Fig. 3. Synthetic geological map of the Sierra Espuña area showing outcrops of the Malaguide units and location of sections.

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Fig. 4. Schematic synthetic stratigraphic columns of the Saladilla Fm, with the location and lithology of studied samples.

In eastern areas (Sierra Espuña; Fig. 3) a tectonic stack of Malaguide units thrusts over the Alpujarride Complex consisting, from top to bottom, of the Perona, Morron de Totana, La Santa, Yéchar and Jaboneros Units (Martín-Martín and Martín-Algarra, 1997; Sanz de Galdeano et al., 2001). The latter three units are exclusively made of Triassic sediments with low-grade metamorphism increasing downwards (Mäkel and Rondeel, 1979). The Morron de Totana Unit is the thickest succession and bears a thin base of Carboniferous sediments below several kilometers of Triassic to lower Miocene sediments (Martín-Martín, 1996; Martín-Martín et al., 1997a, 1997b, 2006a, 2006b). The Malaguide Triassic redbeds have been formally grouped into the Saladilla Formation. It was defined by Roep (1972) in the Eastern Betics, and extended by Martín-Algarra et al. (1995) to the Western Betics and also to the Rifian Ghomarides (see also Zaghloul et al., 2010). This formation (Roep, 1972; Mäkel, 1985; Martín-Algarra et al., 1995; Maate, 1996; Martín-Martín et al., 2006b; Zaghloul et al., 2010) is generally made of quartzose and carbonate clast conglomerates, sandstones and mudstones (siltstones and claystones) with vivid red colors. It also includes calcareous sandstones, dolostones, limestones, marls and gypsum, and is locally intruded by deca- to hectometric bodies of basic sub-volcanic rocks (Soediono, 1971; Geel, 1973; Martín-Algarra, 1987). The Saladilla Fm redbeds are followed by: 1) Rhaetian and Lower Jurassic shallow marine platform carbonates; 2) Pliensbachian–Upper Cretaceous condensed pelagic successions with many hard-grounds; 3) Paleogene shallow marine to basinal limestones and marls; and 4) Upper Oligocene to Lower Miocene deep marine clastics

(Martín-Algarra, 1987; Martín-Martín, 1996; Martín-Martín et al., 1997a, 1997b, 1997c, 2006a; Martín-Algarra et al., 2000, 2009a; Caracuel et al., 2006, among others). Finally, the Malaguide succession is unconformably covered by Burdigalian to Middle Miocene and younger marine to continental clastic deposits. 2.2. The Saladilla Formation The successions of the Saladilla Fm are lithologically quite similar in most Malaguide outcrops and, usually, two transgressive continental to shallow marine megasequences are distinguished (Martín-Algarra et al., 1995). Nevertheless, in the Morron de Totana Unit of Sierra Espuña (Fig. 3), the successions are much thicker, stratigraphically more complete and with a slightly different facies development, especially in its middle and upper part (Fig. 4). There, within beds equivalent to the upper megasequence of other areas, García-Tortosa (2002) and Martín-Martín et al. (2006b) have distinguished smaller sequences made of alternating clastic and carbonate deposits. Similar Triassic successions can be recognized in the lower and middle parts of the underlying La Santa and Yéchar Units of Sierra Espuña (Fig. 3). In addition, the Yéchar Unit shows clear evidence of anchimetamorphism (Mäkel and Rondeel, 1979; Abad et al., 2003; Martín-Martín et al., 2006b). The first megasequence starts with red, medium-to very thicklybedded, usually base-scoured and clast-supported, massive to normallygraded, and sometimes crudely cross-stratified quartzose pebble-to cobble-grade conglomerates (Fig. 5A–D). These alternate with sandstones and mudstones arranged forming predominantly thinning-

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Fig. 5. Main lithofacies of the Saladilla Fm in the Malaga area. (A) Panoramic view of the southern side of the Cerrado de Calderón section, showing the unconformable contact (dashed line) between the Malaguide Palaeozoic and the conglomerate (Cgl) beds of the base of the Saladilla Fm, which evolves upwards to sandstones (Snds). North towards the right. (B) Base of the Middle Triassic redbed succession in the Arroyo de Zapateros section, formed by alternating conglomerates, sandstones and mudstones, unconformably overlying strongly folded, Lower Carboniferous radiolarites. (C) Channeled and megacross-bedded conglomerates of the base of the Cerrado de Calderón succession. (D) Close up of the area squared in (C). (E) View of the southern talus of the Malaga city bypass at the Cerrado de Calderon, close to the Gallery (Street-view photo from GoogleEarth): well bedded, parallel-laminated and cross-laminated sandstones with some pelitic beds and syn-sedimentary normal faults. (F) Alternating carbonate and black clayey beds of the upper detritic and carbonate member of the Saladilla Formation at the Arroyo de los Ángeles. These black clays have provided Carnian pollen with resedimented Carboniferous pollen (for details, see Martín-Algarra et al., 1995). (G) Amalgamated, channeled and wedge-shaped sandstone beds on top of sandsheet deposits. North towards the left; Cerro Bermejo section. (H) Mega-cross-stratified sandbars in medium-grained sandstones with thin mudstone beds. North towards the left; Cerro Bermejo section.

and fining-upwards sequences few meters to a decameter thick, interpreted as channel-fill successions and gravel bars (Roep, 1972; Martín-Algarra et al., 1995; Zaghloul et al., 2010). Laterally, and especially upwards, they evolve to reddish to yellowish, coarse-to fine-grained sandstone bodies interpreted as sandy fluvial facies associations deposited mainly in stream environments related to braided channels and associated bars (Fig. 5E, G–H). These sediments alternate with red mudstones, that locally include pedogenic

carbonate nodules and crusts, and greenish claystones, sometimes bearing gypsum. These fine-grained siliciclastic deposits become progressively thicker upwards and were deposited in flood-plain to coastal mudflat environments (Fig. 6A). Palynological data demonstrate the early Anisian age of this lower part of the succession (Simon and Visscher, 1983; also confirmed by Maate et al., 1993, in the Moroccan Rif near Tetuan). Palaeocurrents testify provenances from areas to the S, SW and SE of the present-day outcrops in most

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Fig. 6. Main lithofacies of the Saladilla Fm in the Sierra Espuña area. (A) Red mudstones and sandstones of the upper part of the lower detritic member, covered by the lower carbonate member; La Santa Unit. (B) to (F) belong to the Morron de Totana Unit. (B) Cross-stratified quartzitic sandstone near to the top of the lower detritic member at Campix; North towards the right. (C) Conglomerates bearing carbonate pebbles and eroding the underlying dolomite bed; upper detritic and carbonate member; the hammer for scale is encircled; Collado Bermejo. (D) Close-up of the conglomerates in (C); note the clast-supported and well rounded texture of the conglomerate, the abundance of carbonate clasts and the presence of rounded clast of intraformational red sandstone. (E) Carbonate beds bearing abundant chert nodules, intercalated within the upper detritic member; the hammer for scale is encircled; Collado Bermejo. (F) Gypsum beds of the uppermost clay and gypsum member; road from Collado Bermejo to Huerta Espuña.

Malaguide areas (Roep, 1972; Mäkel, 1982; Martín-Algarra et al., 1995; see also Figs. 5D and 6B), and from the E or NE in the Rifian Ghomarides (Maate, 1996; Critelli et al., 2008; Zaghloul et al., 2010). The first megasequence culminated with a marine transgression, and Middle Triassic coastal to shallow marine carbonates, with subordinate evaporite beds, were deposited onto the continental redbeds (Fig. 6A). The upper megasequence starts with orange-reddish and varicoloured coastal mudstones and fine-grained sandstones, bearing yellow pebbly sandstones with simple cross-stratification and conglomerate lenses with carbonate clasts (Fig. 6C and D), which were probably deposited in beach environments (Roep, 1972). This megasequence, also includes shallow marine carbonate lenses, sometimes bearing chert (Fig. 6E), and gypsum beds (Fig. 6F) up to several tens of meters thick, deposited in coastal plain sabkhas (Geel, 1973). These beds have provided upper Carnian palynological associations near Málaga (Fig. 5F) and in the Moroccan Rif (Maate et al., 1993; Martín-Algarra et al., 1995). They are covered by Uppermost Triassic and Lower Jurassic shallow marine carbonates.

3. Stratigraphy and facies of the studied Triassic sections Several stratigraphic sections of the Saladilla Fm have been studied both in the Western and Eastern Betics (Figs. 1–6), and some of them were sampled for the purposes of the present study. Three sections were sampled near Malaga (Figs. 2, 5A–H): one along the Malaga city bypass, between the Arroyo de Mayorazgo and the Cerrado de Calderón Gallery; other two in the mountains to the E of Almogía, along the A-3331 road from Malaga to Villanueva de la Concepción close to Cerro Bermejo, and along the Arroyo de Zapateros. In eastern areas, several partial sections along two main transects were sampled in Sierra Espuña, allowing to reconstruct the complete Triassic stratigraphic succession of the Morron de Totana Unit (Figs. 3, 4, 6A–F). 3.1. Western sector (Malaga) The three sections studied near Malaga correspond to the lower part of the Saladilla Fm (mainly the lower megasequence). The Cerrado de Calderón section (Figs. 2 and 4) is more than 250 m thick,

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dips 25° E, and rests on Devonian–Carboniferous siliciclastics. The section is studied along the Malaga bypass roadcut (Figs. 4, 5A and C–D). The Arroyo de Zapateros section (Figs. 2, 4 and 5B) redbeds dip 30° NNW and unconformably overlie Lower Carboniferous radiolarites (O'Dogherty et al., 2000). The Cerro Bermejo section (Figs. 2, 4 and 5G–H) is sub-horizontal to slightly dipping N and near to 100 m thick. The Arroyo de Zapateros section shows more distal facies than those present in the Cerrado de Calderón and also in the Cerro Bermejo sections. In a general way, the Triassic succession from Malaga sector consists at the base of graded quartzose red conglomerates, purple red sandstones and reddish-violet mudstones. Upwardly the succession follows with red, pink, orange, yellow and white, sometimes speckled, thickly-to very thickly-bedded, and fine-to coarse-grained arenites, alternating with vivid red and purple, rarely greenish, mudstones, and with meter- to decimeter-thick intervals made of alternating thinly- to very thinly-bedded fine-grained sandstones and mudstones (Fig. 5E). These deposits (Fig. 5G–H) are interpreted to be deposited by braided rivers, whereas red mudstones deposited in flood-plain to coastal environments. The second transgressive megasequence appears being made of coastal to shallow marine conglomeratic, sandy, pelitic and evaporitic deposits, where upper Carnian pollen has been obtained (Arroyo de los Ángeles: Figs. 2 and 5F; see also Martín-Algarra et al., 1995).

fine-grained reddish to yellowish sandstones, and with thin carbonate and evaporite beds. Upwards, they intercalate some base-scoured and/or channeled mainly clast-supported conglomerate beds (up to 10 m thick), with well-rounded to sub-rounded, quartzose and (predominantly) carbonatic clasts. A thick marine carbonate intercalation develops upwards (marly limestones, limestones and dolostones with cryptalgal lamination) sometimes bearing chert nodules and ribbons. Some intercalated bioclastic packtones–grainstones have provided early Carnian molluscs and echinoderms (Mäkel and Rondeel, 1979; Mäkel, 1985). The youngest beds of the upper detritic member are formed by alternating orange terrigenous mudstones, marls and carbonates with chert (Fig. 6E), including some sandstones, conglomerates and evaporites (Fig. 6C–D) deposited probably in coastal plain to peritidal realms, being the cherty limestones the deepest marine deposits. Above the uppermost carbonate intercalation of Collado Bermejo, the clay and gypsum member makes the top of the Triassic succession (Figs. 3 and 4). It is made of varicoloured and grayish mudstones including thick to very thick beds of laminated gypsum (Fig. 6F), followed by marls alternating with dolostones (sabkha-like environments). Those make the transition to the overlying shallow marine Jurassic carbonates.

3.2. Eastern sector (Sierra Espuña)

Mudstone and sandstone samples were collected along the entire section of the Saladilla Formation in the Malaga and Sierra Espuña areas to cover the entire formation (Fig. 4). The sampling was concentrated in the lower, siliciclastic part of the succession, which correspond to Middle Triassic (Anisian–Ladinian) continental sediments in which the carbonate minerals are least abundant or absent. Nonetheless, to check the evolution of the compositional features upwards in the succession, the Sierra Espuña area was selected for sampling of Upper Triassic terrigenous clastic sediments and associated mudstones (Fig. 4), because: a) the stratigraphic succession is much better preserved, better exposed, and thicker than in the Malaga area; and b) the presence of carbonate components is evident in the field, from the presence of carbonate clasts and grains in both conglomerates and sandstones. We selected 41 medium-to coarse-grained sandstone samples for thin-section preparation and modal analysis. Thin sections were etched and stained for plagioclase and potassium feldspar, and point-counted through the use of a petrographic microscope. From 454 to 581 points were counted for each sample by using the Gazzi–Dickinson method (Ingersoll et al., 1984; Zuffa, 1985). Sandstone grain parameters are defined according to Dickinson (1970), Ingersoll and Suczek (1979), Zuffa (1985), Critelli and Le Pera (1994), and Critelli and Ingersoll (1995). A modal analysis only for the interstitial component, counting 100 points per thin section, was performed by using detailed distinctions of several types of detrital and recrystalized matrix and authigenic components (the latter as pore-filling cements or pore-lining clays and overgrowths of quartz and feldspars). The sandstone samples have been investigated also by SEM-EDX (ESEM Philips Electronics QUANTA 200F with EDX GENESIS 4000). An apatite fission-track dating study was carried out on two sets of apatite grains extracted from two sandstone samples following standard procedures with dense liquids. Fission track dating of apatite is based on the accumulation radiation damage due to the spontaneous nuclear fission of 238U, and it is commonly used to quantify timing and rates of exhumation through temperatures less than 110 °C (for a review of the method see Donelick et al., 2005). In this work, fission track dating was performed according to the external detector method, after irradiation in the atomic reactor TRIGA Mark II at the Oregon State University. Track counting was carried out at DSTGA lab in Bologna, using a microscope Zeiss Axioscope with a total magnification of 1250×, equipped with motorized stage, transmitted and reflected light. Zeta calibration

In Sierra Espuña (Eastern Betics), the succession of the Saladilla Fm is much thicker (around 700 m) and complete than in the Malaga area (Fig. 4). The stratigraphic observations and sampling concentrated in several partial sections along two independent but stratigraphically complementary transects of the Morron de Totana Unit (Casa de la Carrasquilla-Campix and Umbría-Collado Bermejo sections: Fig. 3). The Casa de la Carrasquilla-Campix section (1 in Fig. 3) is about 150 m thick, shows sub-horizontal bedding and corresponds to the lower part of the Triassic redbed succession (first mega-sequence: lower detritic member in Fig. 3). It is followed by thick carbonates (lower carbonate member in Fig. 3), in the same way as observed in other Malaguide units of Sierra Espuña (e.g., La Santa Unit; Fig. 6A). The Umbría-Collado Bermejo section (2 in Fig. 3) dips gently northwards and was measured along a forest track from the Ricardo Codorniu Center to the Collado Bermejo. The succession (upper mega-sequence of the Saladilla Fm) is about 400 m thick from the top of the underlying lower carbonate member to the base of the Jurassic (Fig. 4). In a synthetic way, the succession begins with the lower detritic member unconformably overlying Carboniferous siliciclastics (Almogía Fm). The facies (Fig. 6A–B) are similar to those of the Malaga sector but slightly more distal. In the Sierra Espuña Sector, mudstones are dominant to sandstones (massive to cross-stratified and thickly-to very thickly-bedded: Fig. 6B) and conglomerates. Finally, the redbed succession ends with a mudstone interval, made mainly of claystones with some gypsum and decimeter-thick dolomitic beds near the top, in transition to the overlying dolostone and limestone member (Fig. 3). Triassic successions quite similar to that above described are also present in the other Malaguide tectonic units of Sierra Espuña (La Santa, Yéchar and Jaboneros), as well as in other eastern Malaguide outcrops as Vélez Rubio (Soediono, 1971; Roep, 1972; Geel, 1973). In these units, Middle Triassic to Carnian microfossils (pollen grains, dasycladacean algae, benthic foraminifera and holoturian sclerites) have been found (Mäkel and Rondeel, 1979; Mäkel, 1985), as well as in the carbonates overlying metamorphosed redbeds (Perrone et al., 2006) of the Alpujarride Complex (Los Molinos Unit; Martín-Martín et al., 2006b). The upper megasequence appears over the former carbonates being made up by orange to reddish mudstones that alternate with

4. Sampling and analytical methods

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

was performed according to the procedure recommended by Hurford (1983) using Durango and Fish Canyon Tuff apatites as standards. Neutron fluences were measured counting neutron induced tracks in the Corning glass dosimeters CN-5, placed both at the top and at the bottom of the sample holder to have the maximum control on the neutron fluence gradient. Twenty-three mudstone samples were dried and crushed. Randomlyoriented whole-rock powders were run in the 2–70°2θ interval using a Scintag X1 apparatus. The intensities and diffraction angles of the identified minerals were compared to the database of the International Center for Diffraction Data (ICDD). The mineralogical composition of the b 2 μm grain-size was determined by a thin highly oriented aggregate. The presence of expandable clays was determined after treatment with ethylene glycol at 25 °C for 15 h. Expandability measurements (percent of illite in illite/glycolated-smectite mixed layers) were determined according to Moore and Reynolds (1997). The procedure adopted in this work for the “semiquantitative analysis” is based on the method modified from Moore and Reynolds (1997). Illite crystallinity (IC) measured as full width at half maximum (FWHM) of the 10 Å peak expressed as distance in angular width (Δ° 2θ). Illite crystallinity (IC) value was measured on both air-dried and ethylene-glycol solvated oriented mounts. Illite crystallinity results were calibrated using four interlaboratory Crystallinity Index Standard (CIS) (Warr and Rice, 1994); IC value from this study (x) were transformed into ICCIS data (y) according to the equation y=2.2366x+0.0243 (R2 359=0.98), obtained in laboratory using the international standards (Warr and Rice, 1994). Percentage of illite (%I) and stacking order (Reichweite; R) of the I–S mixed layer was determined on the spectrum of the glycolated specimen according to Moore and Reynolds (1997). Elemental analyses for major and some trace element (Nb, Zr, Y, Sr, Rb, Ba, Ni, Co, Cr, V) concentrations were obtained by X-ray fluorescence spectrometry (Philips PW 1480) on pressed powder disks. X-ray counts were converted into concentrations by a computer program based on the matrix correction method according to Franzini et al. (1972, 1975) and Leoni and Saitta (1976). Total loss on ignition (L.O.I.) was determined, after heating the samples for three hours, at 900 °C. The estimated precision and accuracy for trace element determinations are better than 5%, except for those elements having a concentration of 10 ppm or less (10–15%). 5. Sandstone petrology and detrital modes The obtained raw point-count data of sandstones, including modal data of the detrital and recrystallized matrix and authigenic components, are included in Table 1, whereas the recalculated modal point-count data are shown in Tables 2 (grains) and 3 (interstitial component). The sandstones contain mostly quartz (Figs. 7 and 8A–D) so they range from sublitharenite to quartzarenite sandstones (Fig. 8A). Monocrystalline quartz (Qm) ranges from 94 to 55% of the framework grains; lithic fragments (Lt) are subordinate and range from 41 to 3%. Feldspars (F; both K-feldspar and plagioclases, respectively K and P in Figs. 7 and 8E–F) are minor or absent in many samples (25–0%). Detrital feldspar was partly dissolved and/or replaced by kaolin during diagenesis. Mica and heavy minerals are the main accessory minerals. Lithic fragments are mainly low grade. Felsitic volcanic lithic fragments are subordinately present in a few sandstones. Quartz grains of these sandstones appear as monocrystalline and polycrystalline quartz having tectonic fabric, and as non-foliated polycrystalline quartz, mostly derived from siliceous sedimentary rocks (radiolarian chert) and, rarely, from felsitic volcanic groundmass fragments.

9

of quartz and feldspars (Table 3). The detrital matrix is, generally, a minor component of the whole rock, although it is well preserved in some samples. Due to high compaction and pressure solution processes, the detrital matrix of these sandstones is drastically reduced or recrystallized. Redbed quartzose sandstone has been extensively modified by severe burial history from diagenetic to anchimetamorphic effects, including compaction, intense pressure-solution, precipitation of authigenic minerals, dissolution of framework grains, and severe reduction of porosity. Most detrital grains generally have three or more contacts with neighboring grains, indicating strong mechanical consolidation during compaction. Many mica grains and soft lithic grains (shale, slate and phyllite) and rip-up clasts have undergone ductile deformation (Fig. 8B). Pressure solution features are present in these sandstones, as compenetrated grains, stilolitic and sutured contacts. Triple junctions (Fig. 8A) evidence high distribution of quartz and feldspar overgrowths. Interstitial components were investigated using the petrographic microscope and quantified through point-counting (see Table 3), and using the SEM-EDX. They include quartz, kaolin, Fe-oxides, K-feldspar, albite, dolomite and calcite (see Fig. 9). Quartz cement — Authigenic quartz formed as overgrowths on detrital quartz is the principal cement of these sandstones (Figs. 8C and 9A–B). However, the occurrence of quartz cement is heterogeneous in the studied samples, as related to the presence of continuous clay coat that inhibits quartz overgrowths. Quartz overgrowths postdate the precipitation of dolomite and of Fe-oxides, and, in part, chemical compaction. Carbonate cement — Calcite is one of the most common cements in these sandstones. It appears as poikilotopic crystals, which envelop many grains (Fig. 9C), or as drusy mosaics, which fill the pores between grains. Calcite cement replaces framework grains or other cements in many samples. Fe-oxides — Fe-oxides occur as pore-filling cement (Fig. 9D) and, locally, they have replaced framework grains. They form fibrous crystals, patchy crystals and cements. Fe-oxides predate significant compaction and quartz overgrowth, and seem to postdate some feldspar dissolution. Grain-coating and booklet kaolin — Kaolin (dickite and/or kaolinite) forms felt-like mats of platy crystals arranged tangentially around grains, locally occluding intergranular porosity (Fig. 8C). Graincoating kaolin is commonly mixed with illitic clays (Fig. 8D). XRD analysis indicates that the kaolin mineral is mostly the dickite polytype. Booklet or vermicular kaolin occurs both as pore-filling cement and as a replacement of feldspars (Fig. 9E). Kaolinite is probably formed relatively early in the paragenesis, because it is enclosed by other typical cements as quartz (Figs. 8C and 9E). Feldspar overgrowth — Feldspar overgrowth (Figs. 8E–F and 9G) is a relatively phase of cementation during diagenesis. They are most common on K-feldspars (Figs. 8E and 9G), but they also occur on detrital albite grains (Fig. 8F). Intense dissolution of detrital plagioclase and K-feldspar (Fig. 9H) also takes place during burial. Several minor authigenic phases — Few crystals of authigenic chlorite lamina mainly occur as pore-filling and pore-lining chlorite (Fig. 9F), as such as dolomite includes few remnants of crystals occurring as poikilitic cement with rhombohedral shape. 6. Mineralogical and geochemical composition of mudstones 6.1. Mineralogical results according to stratigraphic sections

5.1. Interstitial composition The interstitial component of these sandstones includes a variety of detrital fine and recrystalized matrix, and authigenic minerals like pore-filling cements or pore-lining clays, as well as overgrowths

The XRD patterns of terrigenous mudstones show that the wholerock mineralogy of the analyzed samples is dominated by clay minerals, associated with significant amounts of quartz and hematite, and minor concentrations of calcite, dolomite and feldspars. The XRD

10

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Table 1 Sandstone raw data. Categories used for sandstone samples point counts and assigned grains in recalculated plots are those of Zuffa (1985, 1987), Critelli and Le Pera (1994), and Critelli and Ingersoll (1995). R.f.=coarse grained rock fragments; NCE=noncarbonate extrabasinal grains; CE=extrabasinal carbonate grains; NCI=noncarbonate intrabasinal grains. Sierra Espuña Lower detritic member

NCE

CE NCI Mx

Cm

TOT

Q Quartz (single crystals) Polycrystalline quartz without tectonic fabric Polycrystalline quartz with tectonic fabric Quartz in volcanic r.f. Quartz in metamorphic r.f. Quartz in plutonic r.f. Quartz in plutonic or gneissic r.f. Quartz in sandstone Calcite replacement on quartz K K-feldspar (single crystals) K-feldspar in volcanic r.f. K-feldspar in metamorphic r.f. K-feldspar in plutonic r.f. K-feldspar in plutonic or gneissic r.f. K-feldspar in sandstone Calcite replacement on k-feldspar P Plagioclase (single crystals) Plagioclase in volcanic r.f. Plagioclase in metamorphic r.f. Plagioclase in plutonic r.f. Plagioclase in plutonic or gneissic r.f. Plagioclase in sandstone Calcite replacement on plagioclase M Micas and e chlorite (single crystals) Mica e clorite in metamorphic r.f. Mica e clorite in plutonic r.f. Mica e clorite in plutonic or gneissic r.f. L Volcanic lithic with felsitic granular texture Metavolcanic lithic Quarzite Slate Phyllite Fine grained schist Impure chert Argillaceous chert Silstone Shale Heavy minerals Heavy minerals (single crystals) Heavy minerals in metamorphic r.f. Opaque minerals Opaque minerals in metamorphic r.f. Opaque minerals in plutonic r.f. Micritic limestone Silty-arenitic limestone Rip-Up clasts Siliciclastic matrix Pseudo- matrix Epi-matrix Carbonate cement (pore-filling) Carbonate cement (patchy-calcite) Phyllosilicate cement Kaolinite cement (pore-filling) Oxid-Fe cement Siliceous cement Quartz overgrowth K-feldspar overgrowth Albite overgrowth Calcite replacement on undeterminated grain Clay and clay grain coats Alterite Undeterminated grain

Upper detritic member

VI-1

VI-9

VI-10

VI-12

VI-13

VI-14

VI-15

VI-16

VI-18

VI-24

VI-27

VI-28

VI-29

VI-30

280 1 2 0 2 0 0 0 0

234 25 17 0 26 0 0 0 0

251 46 40 0 46 0 0 0 0

360 12 4 0 32 0 0 0 0

318 3 3 0 14 0 0 0 0

367 12 2 0 82 0 0 0 0

292 22 4 0 24 0 0 0 0

277 30 2 0 49 0 0 0 0

288 21 13 0 37 0 0 0 0

366 10 14 0 7 0 0 0 4

298 10 19 0 36 0 0 0 9

371 13 16 0 24 0 0 0 8

283 1 4 0 0 0 0 0 0

295 4 6 0 12 0 0 0 1

2 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

61 0 0 0 0 0 0

111 0 0 0 0 0 0

49 0 1 0 2 0 0

9 0 0 0 0 0 0

8 0 0 0 0 0 0

0 0 0 0 0 0 0

7 0 0 0 0 0 1

2 0 0 0 0 0 0

23 0 0 0 0 0 0

2 0 0 0 0 0 0

17 0 0 0 0 0 0

4 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

5 0 0 0 0 0 0

2 0 0 0 0 0 0

11 0 0 0 0 0 0

11 0 0 0 0 0 0

2 0 1 0 0 0 0

4 0 0 0 0 0 1

4 0 1 0 0 0 0

11 0 1 0 0 0 2

0 0 0 0 0 0 0

6 0 0 0 0 0 0

13 0 0 0

4 0 0 0

0 0 0 0

2 0 0 0

14 0 0 0

2 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

3 0 0 0

1 0 0 0

0 0 0 5 5 0 0 0 0 0

0 0 0 20 35 0 5 0 0 0

0 0 0 2 8 0 48 1 0 0

1 0 0 3 4 0 2 1 0 0

0 0 0 6 2 0 0 0 0 0

0 0 0 3 5 0 3 1 0 0

0 0 0 0 9 1 0 0 0 0

0 0 0 0 16 0 0 0 0 0

0 0 0 24 33 0 7 0 0 0

0 0 0 1 13 0 3 0 0 0

0 0 0 0 4 0 11 0 0 0

0 0 0 1 4 0 4 0 0 0

0 0 0 6 11 0 0 0 0 0

1 0 0 1 6 0 2 0 0 0

2 0 5 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 156 3 1 0 481

1 0 1 0 0 0 0 0 46 0 0 0 0 0 25 33 0 0 0 0 0 7 5 0 484

0 0 0 0 0 0 0 0 1 0 0 0 0 0 12 0 0 4 0 0 0 6 4 0 469

3 0 0 0 0 0 0 0 0 0 0 0 0 0 0 3 0 11 6 1 0 7 0 0 513

5 0 4 0 0 0 0 0 2 0 0 0 0 0 0 7 0 1 9 0 0 47 1 0 552

1 1 2 0 0 0 0 1 2 0 0 0 0 0 0 0 0 3 0 2 0 4 1 0 548

5 0 0 0 0 0 0 0 10 0 0 1 12 1 0 34 0 55 2 2 6 0 0 0 500

2 0 0 0 0 0 0 0 16 0 0 0 0 0 0 17 2 65 2 11 0 0 0 0 508

0 0 0 0 0 0 0 1 13 0 0 0 0 0 33 33 0 1 0 0 0 18 1 0 526

2 0 3 0 0 0 0 0 5 0 0 1 6 0 0 10 0 0 0 0 73 4 2 0 537

0 0 0 0 0 0 0 0 1 0 0 1 0 0 0 0 0 0 0 0 155 0 1 0 552

3 0 2 0 0 0 0 16 5 0 0 11 2 0 0 10 0 2 1 1 43 7 0 0 581

1 0 3 0 0 0 0 0 80 0 0 0 0 0 0 32 0 0 0 0 124 2 1 0 553

2 1 2 0 0 0 0 0 0 0 0 1 1 0 0 0 0 1 0 0 175 0 0 0 535

Note: Categories used for sandstone samples point counts and assigned grains in recalculated plots are those of Zuffa (1985, 1987), Critelli and Le Pera (1994), and Critelli and Ingersoll (1995). R.f. = coarse grained rock fragments; NCE = noncarbonate extrabasinal grains; CE = extrabasinal carbonate grains; NCI = noncarbonate intrabasinal grains.

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

11

Cerro Bermejo VII-1

VII-2

VII-4

VII-5

VII-9

VII-13

VII-15

VII-16

VII-17

VII-18

VII-19

VII-20

VII-20''

VII-22

VII-23

VII-25

VII-26

VII-27

VII-28

196 21 3 0 63 0 0 0 0

181 21 5 0 57 0 0 0 0

154 77 16 0 67 0 0 0 2

293 22 18 0 54 0 0 0 0

191 14 15 0 83 0 0 0 3

382 15 0 0 69 0 0 0 0

348 20 4 0 84 0 0 0 0

273 19 20 0 49 0 0 0 1

367 8 3 0 93 0 0 0 0

347 12 4 0 89 0 0 0 0

183 33 30 0 94 0 0 0 1

243 21 24 0 100 0 0 0 2

164 20 26 0 83 0 0 0 2

311 14 3 0 65 0 0 0 0

125 84 32 0 168 0 0 0 0

258 65 14 0 137 0 0 0 0

187 85 17 0 132 0 0 0 0

295 37 8 0 55 0 0 0 0

285 10 7 0 96 0 0 0 0

4 0 0 0 0 0 0

2 0 2 0 0 0 0

0 0 0 0 1 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

35 0 0 0 0 0 0

33 0 0 0 0 0 0

0 0 0 0 0 0 0

16 0 0 0 0 0 0

9 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

1 0 0 0 0 0 0

7 0 0 0 0 0 0

2 0 0 0 0 0 0

4 0 0 0 0 0 0

1 0 0 0 0 0 0

9 0 0 0 0 0 0

8 0 0 0 0 0 0

18 0 21 1 0 0 0

32 0 16 0 0 0 0

0 0 0 0 0 0 0

2 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

16 0 0 0 0 0 0

10 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 0 0 0 0 0

0 0 1 0 0 0 0

23 0 1 0 0 0 0

2 0 0 0 0 0 0

4 0 1 0 0 0 0

2 0 1 0 0 0 0

20 0 0 0 0 0 0

15 0 3 0 0 0 0

32 0 2 4

45 0 0 2

1 0 0 0

0 0 0 0

1 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

2 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

0 0 0 0

2 0 0 0

2 0 0 0

3 0 0 0

2 0 0 0

3 0 0 3

1 1 0 7 29 3 5 0 0 0

3 0 0 8 38 0 4 0 0 0

0 0 0 1 29 0 16 7 0 1

0 0 0 4 6 0 10 0 0 0

0 0 0 25 14 0 5 1 0 0

0 0 0 1 5 0 0 0 0 0

0 0 0 0 7 0 2 0 0 0

0 0 0 15 19 0 7 0 0 0

0 0 0 1 6 0 0 0 0 0

0 0 0 0 6 0 0 0 0 0

0 0 0 8 3 0 9 0 0 0

1 1 0 13 11 0 7 0 1 0

0 0 2 10 15 0 5 0 0 0

0 0 0 1 11 0 1 0 0 0

0 2 7 2 9 0 22 0 0 0

0 0 0 1 20 0 12 2 0 0

1 0 0 0 11 0 39 1 0 0

0 0 0 2 28 0 2 0 0 0

0 1 0 2 13 0 0 0 0 0

0 0 4 0 0 0 0 0 0 0 0 0 0 0 0 2 0 0 0 0 4 50 6 0 477

1 0 5 0 0 0 0 0 1 0 0 0 0 0 0 2 0 0 0 0 2 31 3 0 461

0 0 0 0 0 0 0 0 8 0 0 6 7 0 34 25 0 33 0 0 20 2 1 1 509

0 0 1 0 0 0 0 0 6 0 0 0 0 0 33 15 0 6 0 0 15 6 1 0 492

1 0 0 0 0 0 0 24 13 0 0 0 10 0 27 13 0 9 0 0 102 11 3 0 565

3 0 0 0 0 0 0 2 1 0 0 0 0 0 0 3 0 26 7 0 0 0 0 0 549

0 0 0 0 0 0 0 2 0 0 0 0 0 1 0 30 0 1 4 0 0 1 0 0 537

1 0 2 0 0 0 0 1 3 0 0 0 0 0 56 14 0 0 0 0 47 5 3 0 535

3 0 0 0 0 0 0 1 1 0 0 0 0 0 0 13 0 16 0 7 0 1 0 0 552

2 0 0 0 0 0 0 0 3 0 0 0 0 0 0 34 0 28 0 9 0 8 0 0 563

0 0 0 0 0 0 0 20 1 0 0 0 0 0 24 36 0 3 0 0 22 12 2 0 481

2 0 0 0 0 0 0 6 4 0 0 0 1 0 43 22 0 15 0 0 31 18 3 0 569

0 0 0 0 0 0 0 17 3 0 0 0 4 0 51 20 0 7 0 0 40 10 0 0 481

0 0 0 0 0 0 0 0 2 0 0 0 0 0 0 20 1 52 3 16 0 1 0 0 532

0 0 2 0 0 0 0 5 2 0 0 0 0 1 0 0 1 2 0 1 0 13 0 0 484

0 0 0 0 0 0 0 1 2 0 0 0 0 0 14 2 0 20 0 4 0 9 0 0 572

1 0 0 0 0 0 0 9 0 0 0 0 0 1 0 0 0 7 0 0 0 4 0 0 502

4 0 3 0 0 0 0 0 9 0 0 0 0 0 0 32 7 7 1 3 0 1 1 0 526

1 0 4 0 0 0 0 0 11 0 6 0 0 0 0 6 0 40 0 2 0 1 1 0 518

(continued on next page)

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F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Table 1 (continued) Cerrado de Calderón

NCE

CE NCI Mx

Cm

TOT

Q Quartz (single crystals) Polycrystalline quartz without tectonic fabric Polycrystalline quartz with tectonic fabric Quartz in volcanic r.f. Quartz in metamorphic r.f. Quartz in plutonic r.f. Quartz in plutonic or gneissic r.f. Quartz in sandstone Calcite replacement on quartz K K-feldspar (single crystals) K-feldspar in volcanic r.f. K-feldspar in metamorphic r.f. K-feldspar in plutonic r.f. K-feldspar in plutonic or gneissic r.f. K-feldspar in sandstone Calcite replacement on k-feldspar P Plagioclase (single crystals) Plagioclase in volcanic r.f. Plagioclase in metamorphic r.f. Plagioclase in plutonic r.f. Plagioclase in plutonic or gneissic r.f. Plagioclase in sandstone Calcite replacement on plagioclase M Micas and e chlorite (single crystals) Mica e clorite in metamorphic r.f. Mica e clorite in plutonic r.f. Mica e clorite in plutonic or gneissic r.f. L Volcanic lithic with felsitic granular texture Metavolcanic lithic Quarzite Slate Phyllite Fine grained schist Impure chert Argillaceous chert Silstone Shale Heavy minerals Heavy minerals (single crystals) Heavy minerals in metamorphic r.f. Opaque minerals Opaque minerals in metamorphic r.f. Opaque minerals in plutonic r.f. Micritic limestone Silty-arenitic limestone Rip-Up clasts Siliciclastic matrix Pseudo- matrix Epi-matrix Carbonate cement (pore-filling) Carbonate cement (patchy-calcite) Phyllosilicate cement Kaolinite cement (pore-filling) Oxid-Fe cement Siliceous cement Quartz overgrowth K-feldspar overgrowth Albite overgrowth Calcite replacement on undeterminated grain Clay and clay grain coats Alterite Undeterminated grain

VIII-5

VIII-7

VIII-10

VIII-12

VIII-14

VIII-17

VIII-19

VIII-22

325 12 4 0 54 0 0 0 0

209 47 19 4 116 0 0 0 19

150 74 20 3 53 0 0 0 13

303 9 5 0 60 0 0 0 0

220 27 18 0 71 0 0 0 1

359 4 3 0 37 0 0 0 0

283 18 15 0 37 0 0 0 7

347 5 5 0 41 0 0 0 0

32 0 0 0 0 0 0

15 0 0 0 0 0 0

11 0 0 0 0 0 0

7 0 0 0 1 0 0

4 0 0 0 0 0 0

36 0 0 0 0 0 0

38 0 0 0 0 0 0

45 0 0 0 0 0 0

1 0 0 0 0 0 0

20 0 2 1 0 0 0

7 0 0 0 0 0 0

9 0 0 0 0 0 0

0 0 0 0 0 0 0

1 0 0 0 0 0 0

0 0 0 0 0 0 0

2 0 0 0 0 0 0

3 0 0 0

0 0 0 0

0 0 0 1

1 0 0 0

1 0 0 0

9 0 0 0

1 0 0 0

7 0 0 0

0 0 0 2 3 0 0 0 0 0

7 2 0 2 14 0 4 0 0 0

5 2 0 4 38 0 18 0 1 0

0 0 0 2 2 0 0 0 0 0

2 0 0 4 28 0 18 1 0 0

0 0 0 3 7 0 0 0 0 0

0 1 0 5 12 0 2 0 0 0

0 0 0 2 15 0 0 0 0 0

2 0 0 0 0 0 0 7 2 0 0 0 0 0 0 13 0 56 7 0 0 2 0 0 525

0 0 0 0 0 0 0 0 1 0 0 4 11 0 0 2 0 5 0 7 42 2 1 0 556

0 0 1 0 0 0 1 0 1 0 0 10 11 0 0 4 0 2 0 0 91 1 1 0 523

3 0 0 0 0 0 0 7 1 0 0 0 0 0 0 15 1 18 1 8 0 0 1 0 454

3 0 0 0 0 1 0 2 1 0 0 0 6 0 0 6 0 2 4 1 105 1 0 0 527

11 0 3 0 0 0 0 0 6 0 0 0 0 0 0 13 0 34 8 1 0 0 0 0 535

1 0 0 0 0 0 0 0 3 0 0 6 13 0 0 2 0 1 8 0 91 0 4 0 548

4 0 1 0 0 0 0 0 1 0 0 0 0 0 0 10 0 5 1 1 0 23 1 0 516

patterns of the Upper Triassic sediments (upper detritic member) of the Saladilla Fm at Sierra Espuña (Fig. 4) show intense calcite and dolomite peaks. The mudstones collected in the lower detritic members (Middle Triassic) of the western areas (Cerrado de Calderón, Cerro Bermejo and

Arroyo de Zapateros sections; Figs. 2 and 4) display XRD patterns characterized by minor or trace amounts of calcite and dolomite. The b2 μm grain-size fraction is composed by illite (ranging from 64% to 80%) prevailing on illite–smectite mixed layers (I–S) (ranging

Table 2 Recalculated modal point count data. X= mean; SD = standard deviation. Qm = monocrystalline quartz, F = feldspars (K + P), Lt= lithic grains; K = K-feldspar, P = plagioclase; Qp = polycrystalline quartz, Lvm = volcanic and metavolcanic, Lsm = sedimentary and metasedimentary. Age

Sample

Sierra Espuña

Upper Triassic

VI-30 VI-29 VI-28 VI-27 VI-24 VI-18 VI-16 VI-15 VI-14 VI-13 VI-12 VI-10 VI-9 VI-2 VII-1 VII-2 VII-4 VII-5 VII-9 VII-13 VII-15 VII-16 VII-17 VII-18 VII-19 VII-20 VII-21 VII-22 VII-23 VII-25 VII-26 VII-27 VII-28 VIII-5 VIII-7 VIII-10 VIII-12 VIII-14 VIII-17 VIII-19 VIII-22 X SD

Middle Triassic

Cerro Bermejo

Middle Triassic

Cerrado de Calderón

Middle Triassic

%

%

%

%

%

%

%

Qm

F

Lt

Qt

F

L

Qm

K

P

Qp

Lvm

Lsm

Lm

Lv

Ls

Rg

Rv

Rm

Rg

Rs

Rm

88 92 84 87 87 76 83 85 85 72 82 67 72 94 69 64 60 85 79 89 87 80 90 91 77 81 76 86 64 76 67 77 87 88 72 55 91 74 88 78 84 79 ±9

7 1 8 2 3 1 5 5 10 25 13 0 0 2 12 14 0 0 0 7 7 0 6 4 0 0 1 7 1 2 1 6 6 8 8 5 4 1 8 9 10 5 ±5

6 7 8 11 10 23 12 10 5 3 6 33 28 4 19 21 40 15 21 4 7 20 4 5 23 19 24 7 35 22 32 17 8 5 20 41 5 25 4 13 6 16 ±11

91 94 91 97 94 86 91 92 88 73 86 98 85 95 77 73 91 97 89 92 92 92 92 95 97 94 91 90 95 94 97 87 90 91 87 83 95 90 90 87 86 90 ±6

7 1 8 2 3 1 5 5 10 25 13 0 0 2 12 14 0 0 0 7 7 0 6 4 0 0 1 7 1 2 1 6 6 8 8 5 4 1 8 9 10 5 ±5

2 6 1 1 3 13 4 3 2 2 2 2 15 3 11 13 8 2 11 1 1 8 1 1 3 6 8 3 4 4 3 7 4 1 5 13 1 9 2 4 4 5 ±4

93 99 92 98 97 99 94 94 89 74 87 100 100 98 85 82 100 99 100 93 93 100 93 96 100 100 99 92 99 98 99 92 94 92 90 92 96 99 91 90 89 94 ±6

5 1 5 1 2 0 2 3 10 25 13 0 0 1 1 1 0 0 0 7 7 0 3 2 0 0 0 2 1 1 0 2 2 8 4 5 2 1 8 10 10 4 ±5

2 0 3 1 1 1 3 3 0 1 0 0 0 1 13 17 0 1 0 0 0 0 3 2 0 0 0 6 1 1 1 5 4 0 6 3 2 0 0 0 0 2 ±3

60 23 87 91 66 42 67 72 69 43 70 93 46 23 41 38 79 83 47 71 79 58 61 73 87 66 67 60 91 82 92 61 52 76 74 69 78 65 41 66 37 65 ±18

5 0 0 0 0 0 0 0 0 0 4 0 0 0 3 4 0 0 0 0 0 0 0 0 0 3 0 0 1 0 1 0 3 0 9 4 0 2 0 2 0 1 ±2

35 77 13 9 34 58 33 28 31 57 26 7 54 77 56 58 21 17 53 29 21 43 39 27 13 32 33 40 7 18 7 39 45 24 17 27 22 33 59 32 63 34 ±19

70 100 56 27 82 89 100 100 67 100 64 17 92 100 87 87 56 50 87 100 78 83 100 100 55 74 84 92 48 60 21 94 100 100 62 65 100 60 100 90 100 79 ±23

10 0 0 0 0 0 0 0 0 0 9 0 0 0 2 6 0 0 0 0 0 0 0 0 0 3 0 0 0 0 2 0 0 0 24 7 0 4 0 0 0 2 ±4

20 0 44 73 18 11 0 0 33 0 27 83 8 0 11 8 44 50 13 0 22 17 0 0 45 24 16 8 52 40 77 6 0 0 14 28 0 36 0 10 0 19 ±23

0 0 0 0 0 0 0 0 2 0 0 0 0 0 5 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 1 1 1 0 0 0 0 0 ±1

4 0 0 0 0 0 0 0 0 0 2 0 0 0 1 2 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 6 4 0 2 0 0 0 0 ±1

96 100 100 100 100 100 100 100 98 100 98 100 100 100 95 97 99 100 100 100 100 100 100 100 100 99 100 100 100 100 100 100 98 100 94 95 99 98 100 100 100 99 ±2

0 0 0 0 0 0 0 0 2 0 0 0 0 0 4 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 1 1 1 0 0 0 0 0 ±1

8 0 9 18 9 6 0 0 4 0 6 32 5 0 3 3 12 10 4 0 2 6 0 0 6 5 4 1 7 6 15 2 0 0 2 10 0 13 0 3 0 5 ±6

92 100 91 82 91 94 100 100 95 100 94 68 95 100 92 96 87 90 96 100 98 94 100 100 94 95 96 99 93 94 85 98 98 100 97 89 99 87 100 97 100 95 ±6

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Section

Note: X = mean; SD = standard deviation. Qm = monocrystalline quartz, Qp = polycrystalline quartz, F = feldspars (K + P), K = K-feldspar, P = plagioclase; Lt = lithic grains; Lm = metamorphic, Lv = volcanic, and Ls = sedimentary lithic grains; Lvm = volcanic and metavolcanic, Lsm = sedimentary and metasedimentary lithic grains; Rg = phaneritic plutonic rock fragments; Rm = coarse and fine grained metamorphic rock fragments; Rv = coarse and fine grained volcanic rock fragments; Rs = coarse and fine grained sedimentary rock fragments.

13

14

Table 3 Point counting analysis of the interstitial component of Malaguide sandstones. Sierra Espuña VI-9

VI-10

VI-12

VI-13

VI-14

VI-15

VI-16

VI-18

VI-24

VI-27

VI-28

VI-29

VI-30

VII-1

VII-2

VII-4

VII-5

VII-9

VII-13

VII-15

0 19 17 0 0 0 0 0 0 0 0 63 1 0 0

38 4 12 0 4 0 16 0 0 0 0 0 26 0 0

0 4 6 0 37 0 37 0 0 0 0 0 16 0 0

19 2 14 5 26 9 0 0 0 0 0 0 2 18 5

18 15 24 5 5 0 0 0 0 0 0 0 11 12 10

0 8 21 9 21 14 0 0 0 0 0 0 3 15 9

21 0 2 2 24 3 0 0 0 3 2 11 22 3 7

19 2 3 6 31 4 0 0 0 0 0 0 6 13 16

30 2 17 0 5 0 26 0 1 0 0 0 19 0 0

23 4 11 0 5 0 0 0 0 6 1 28 15 4 3

16 2 6 0 11 0 0 0 0 7 1 53 1 2 1

37 1 10 0 1 0 0 0 0 16 3 13 3 9 7

37 11 2 0 0 0 0 0 0 0 0 49 1 0 0

3 9 12 0 10 0 0 0 0 2 0 44 5 12 3

4 7 43 16 1 0 0 25 0 0 0 3 1 0 0

1 23 47 16 0 1 0 8 0 0 0 1 3 0 0

9 2 9 4 20 2 30 0 0 6 2 9 7 0 0

16 9 9 3 11 1 19 0 0 10 0 8 14 0 0

28 7 5 2 2 0 18 0 0 7 2 24 5 0 0

1 6 3 4 47 4 0 0 0 0 0 0 2 26 7

9 2 13 2 20 0 0 0 0 0 0 0 29 23 2

Table 3 (continued) Cerro Bermejo

Ortomatrix Pseudo-matrix Epi-matrix Protomatrix (Clay grain coats) Quartz overgrowth Siliceous cement Kaolinite cement (pore-filling) Chlorite cement (pore-filling) Clay e clay grain coats Carbonate cement (pore-filling) Carbonate cement (patchy-calcite) Calcite replacement on undeterminated grain Oxid-Fe cement K-feldspar overgrowth Albite overgrowth

Cerrado de Calderón

VII-16

VII-17

VII-18

VII-19

VII-20

VII-21

VII-22

VII-23

VII-25

VII-26

VII-27

VII-28

VIII-5

VIII-7

VIII-10

VIII-12

VIII-14

VIII-17

VIII-19

VIII-22

18 12 7 1 2 0 29 0 0 1 0 15 15 0 0

6 3 2 0 38 2 0 0 0 0 0 0 10 20 19

9 7 2 7 33 4 0 0 0 0 0 0 16 13 9

18 3 13 7 10 1 19 0 0 3 1 11 14 0 0

21 0 8 2 18 1 30 0 0 0 0 10 10 0 0

16 0 1 2 17 2 36 0 0 5 1 9 11 0 0

5 5 3 1 40 12 0 0 0 0 0 0 6 14 14

0 12 20 39 13 4 0 0 7 0 0 0 0 3 2

5 6 10 17 30 7 0 0 20 0 0 0 1 2 2

7 10 23 28 18 7 0 3 0 0 0 0 0 2 2

23 8 7 6 23 8 0 0 0 0 0 0 6 7 12

4 11 12 5 30 12 0 0 0 0 0 0 7 8 11

9 7 1 23 32 0 0 0 0 0 0 0 6 16 6

5 2 7 0 9 0 0 0 0 16 4 23 3 14 17

3 3 6 6 3 0 0 0 0 7 3 45 15 8 1

13 4 0 22 51 3 0 0 0 0 0 0 4 3 0

5 3 3 2 1 0 0 0 0 6 0 59 16 2 3

3 14 0 35 29 0 0 0 0 0 0 0 5 11 3

3 9 4 0 5 0 0 0 0 7 1 38 6 20 7

7 23 27 20 13 0 0 0 0 0 0 0 1 9 0

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Ortomatrix Pseudo-matrix Epi-matrix Protomatrix (Clay grain coats) Quartz overgrowth Siliceous cement Kaolinite cement (pore-filling) Chlorite cement (pore-filling) Clay e clay grain coats Carbonate cement (pore-filling) Carbonate cement (patchy-calcite) Calcite replacement on undeterminated grain Oxid-Fe cement K-feldspar overgrowth Albite overgrowth

Cerro Bermejo

VI-2

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

15

Fig. 7. Qt–F–L, Qm–F–Lt, Qm–K–P, Rg–Rs–Rm, Qp–Lvm–Lsm and Lm–Lv–Ls triangular plots (from Dickinson, 1970; Ingersoll and Suczek, 1979; Ingersoll et al., 1984; Zuffa, 1985; Critelli and Le Pera, 1994; Critelli and Ingersoll, 1995) for Malaguide redbed sandstones of the Betic Cordillera. Qt (quartz grains), F (feldspars) and L (aphanitic lithic fragments); Qm (monocrystalline quartz), F (feldspars) and Lt (total lithic fragments); Qm (monocrystalline quartz), K (K-feldspar) and P (plagioclase); Rg (plutonic rock fragments), Rs (sedimentary rock fragments) and Rm (metamorphic rock fragments); Qp (polycrystalline quartz), Lvm (volcanic and metavolcanic lithic fragments) and Lsm (sedimentary and metasedimentary lithic fragments); Lm (metamorphic), Lv (volcanic) and Ls (sedimentary) lithic fragments.

from 17% to 32%), with minor amounts of kaolinite and chlorite (ranging from trace to 10%). In the studied samples, the most significant clay mineral variations are principally associated with the % of I–S mixed layers and with the content of illitic layers in I–S. The XRD traces of the b 2 μm fraction show a little decrease in the proportion of I–S mixed layers at the base of the stratigraphic succession. Furthermore, the detailed low-angle scans of the air dried and ethylene-glycol solvated b2 μm fraction specimens of the lower (Middle Triassic) samples show small peak shifts of I–S, implying that the proportion of smectite-like layers in interstratified I–S is relatively low. In the Upper Triassic samples, the peak shifts of I–S are more evident, consequent to ethylene-glycol solvation. This is probably related to the diagenetic evolution recorded by the samples (e.g. Perri, 2008). 6.2. Major and trace element variations The mudstone elemental concentrations are given in Table 4. The elemental distributions normalized to Post-Archean Australian Shale (PAAS; Taylor and McLennan, 1985) are shown in Fig. 10. The mudstones are characterized by narrow compositional changes for Si, Ti, Al, Fe, and K, which have concentrations close to those of PAAS (Fig. 10A). Sodium, manganese and phosphorous are strongly depleted relatively to PAAS, whereas Ca shows high variability in concentration ranging from 0.24 (FP60) to 16.19 wt.% (FP71). The observed Na depletion is likely due to the burial history of these samples, which promoted

the formation of K-rich, mica-like clay minerals. The higher Ca concentrations are related to the carbonate minerals present in some Upper Triassic samples of the Sierra Espuña section. Mg is enriched relatively to PAAS, due to the dolomite occurring in the same samples. The range of all trace elements but thorium appears to overlap the PAAS (Fig. 10B). Furthermore Rb, shows positive linear correlation with K, suggesting that this element is mostly controlled by the mica-like clay minerals. Samples FP70 and FP71, collected along the Sierra Espuña section nearby to carbonatic and evaporitic beds have the highest Ba content (2318 and 6087 ppm), well above the PAAS content (650 ppm). All the transition metals (V, Cr, Co, and Ni) show positive linear correlations with Al2O3, Fe2O3 and K2O. The high field strength elements (HFSE) have concentrations quite similar to those of PAAS. The concentrations of the transition metals and the ratios among Cr/V and Y/Ni are also close to the PAAS. The studied sediments show the La/Co (average= 2.67), La/Cr (average=0.54) and La/Ni (average=1.23) ratios quite similar to the PAAS (La/Co=1.65; La/Cr=0.35; La/Ni=0.69) and upper continental crust (UCC: McLennan et al., 2006; La/Co=3.00; La/Cr=0.86; La/Ni=1.50) ratios. 7. Fission-track study For the apatite fission-track study, two sandstone samples were selected from the top of the Cerro Bermejo section (Middle Triassic part of the succession in the Malaga area (sample FP-50ar) and from

16

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Fig. 8. Types of interstitial components in sandstone (crossed nicols view). (A): Pore-filling cement constitutes by authigenic clay minerals (10×); (B): pseudomatrix composed of soft grains like rip-up clast (20×); (C): pore-filling kaolinite cement (Kaol) and quartz overgrowth (Q-o) (20×); (D): mixed layers of authigenic clay minerals of illite and kaolinite (Kaol) (20×); (E): K-feldspar overgrowth (K-o) on detrital K-feldspar grain (K) (20×); and (F): albite overgrowth (P-o) on detrital plagioclase grain (P) (20×).

the base of the Sierra Espuña succession (sample FP87ar), respectively (Fig. 4)). Sample FP50ar yielded only 11 apatite crystals that could be dated (Table 5). Therefore, the error in the age calculation is quite high (about 40% for 1σ error), although the Chi-square test is passed and no dispersion of single grain ages is present. The high error is due also to the very low uranium content (5.2 ppm). The low number of etched tracks makes impossible also any measure of their length. However, we can argue that these data suggest a thermal history very similar to what was observed in sample FP87ar (Table 5). Twenty grains have been dated in sandstone sample FP87ar (Table 5). The Chi-square test is passed, thus demonstrating a very low dispersion in grain age distribution. The calculated central age (15.6 Ma) is much younger than the stratigraphic age, therefore indicating that temperatures reached during post-depositional burial were higher than total resetting temperature (about 130 °C). The calculated ages should reflect the timing of cooling below the ca. 110 °C isotherm. Unfortunately, in the samples here analyzed a statistically meaningful amount of track lengths could not be measured and, consequently, it cannot be demonstrated that no significant annealing has occurred in post-closure time (Green et al., 1989). 8. Subsidence analysis and compaction To compare the geological evolution of the Malaguide successions with the changes in clay mineralogy of the studied redbeds (see later

sections), a study of the basin subsidence and compaction of the Sierra Espuña succession (Morron de Totana Unit) has been carried out. The Morron de Totana Unit was selected because of its Triassic and younger stratigraphic succession is the most complete and thickest, and the less disturbed tectonically, of the Malaguide Complex (see Martín-Martín et al., 2006b, and Fig. 4, for details of the stratigraphic column of the Triassic; Caracuel et al., 2006; Martín-Martín et al., 2006a, for the Jurassic–Cretaceous stratigraphy; and Martín-Martín, 1996; Serra-Kiel et al., 1996/1998; Martín-Martín et al., 1997a, 1997b, 1997c, 1998, for the Tertiary stratigraphy). Subsidence analysis has been performed using standard procedures (Watts and Ryan, 1976; Steckler and Watts, 1978; Sclater and Christie, 1980; Xie and Heller, 2009) using the program “Loadcap” by “Geostru Software” licensed to the Alicante University (reference no. G38RJ2). This study takes into account the load due to deposition of each sedimentary sequence of the basin infilling, the stacked tectonic units and also the load due to the mean water column, calculated from paleobathymetric interpretation (see Appendix A for details on the subsidence analysis and compaction study). The results obtained are shown in Table 6 and in Fig. 11, indicating that the measured thicknesses in the field (Fig. 11A) increase in different ways according to the original lithologies when decompaction is performed (Fig. 11B). So, the depth of the Palaeozoic–Triassic boundary, which is covered by sediments reaching a maximum total thickness of 3025 m (see Fig. 11A), becomes located at 5370 m deep

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

17

Fig. 9. SEM photomicrographs of the main diagenetic phases in the redbed sandstones. (A): Polygonal triple junction of quartz overgrowth; (B): quartz overgrowth (Q-o) on detrital quartz grain (Q); (C): mica detrital grain (M) in pore-filling calcite cement; (D): Fe-oxide cement (Fe-ox); (E): kaolinite cement evolved by a quartz cement crystal; (F): chlorite cement (Chl) occurs as pore-lining; (G): authigenic K-feldspar overgrowth (K-o); and (H): corroded, detrital K-feldspar grain (K).

after decompaction (see Fig. 11B). Also, the sedimentary sequences made of soft sediments (silt, clay, marl, sand and gypsum) become highly compacted. This is particularly evident in the case of the Triassic package, which reaches a thickness of 1775 m after decompaction although it has a maximum measured thickness of around 700 m (Fig. 4). In the same way, the present thickness of the Paleocene–Eocene

package (275 m) becomes 1092 m after decompaction. In the former case, the compaction was very important since the lithologies in this sequence are very soft and suffered the loading of the very thick and comparatively harder Oligocene sequence, and also the tectonic overload due to thrusting of the Perona Unit (Martín-Martín and Martín-Algarra, 2002). The thickness of both the Oligocene plus the Perona Unit on the

18

Table 4 Major, trace element and ratios distribution of the mudstone redbed samples. Age

Middle Triassic

Sections

Cerro Bermejo

Samples

FP46

FP47 55.74 0.76 15.72 5.32 0.08 4.69 4.50 0.26 3.91 0.07 8.92 99.97

Trace elements (ppm) Sc 16 V 163 Cr 109 Co 25 Ni 57 Rb 208 Sr 107 Y 39 Zr 189 Nb 21 Ba 758 La 68 Ce 112 Th 11

13 102 77 18 39 132 123 31 247 16 394 45 89 11

Ratios La/Sc Th/Sc La/Cr La/Ni CIW CIA PIA

3.46 0.86 0.58 1.15 N.D N.D N.D

4.15 0.68 0.62 1.19 92.95 70.70 89.72

FP48 51.11 0.96 21.95 9.68 0.03 2.73 0.32 0.23 6.49 0.13 6.34 99.97

18 171 134 29 50 206 422 39 214 18 601 70 138 9

3.80 0.50 0.52 1.40 96.70 66.79 94.02

FP49 52.60 1.07 21.60 9.35 0.02 2.44 0.29 0.22 6.56 0.18 5.65 99.98

25 169 149 21 40 213 429 56 332 23 803 88 183 8

3.52 0.31 0.59 2.20 96.88 66.32 94.21

FP50 53.96 0.96 21.09 8.55 0.03 2.69 0.33 0.22 6.29 0.16 5.43 99.71

22 159 121 26 49 203 419 42 231 21 623 69 145 9

3.21 0.44 0.57 1.41 96.57 66.55 93.75

FP51 55.45 0.84 19.23 8.58 0.03 3.51 0.93 0.58 4.23 0.14 6.10 99.62

19 133 107 27 47 229 95 31 174 23 442 54 82 8

2.84 0.42 0.50 1.15 90.34 68.90 85.98

FP52 54.06 0.76 19.55 8.12 0.03 4.41 0.71 0.48 3.88 0.14 7.45 99.59

18 132 102 30 47 238 83 32 187 21 394 45 89 8

2.46 0.45 0.44 0.96 92.34 71.75 89.26

FP53 56.01 0.78 19.40 6.97 0.07 3.33 2.31 0.58 3.97 0.11 6.25 99.78

18 125 92 21 41 178 79 38 218 20 509 61 92 9

3.49 0.50 0.66 1.49 87.19 68.15 82.22

FP54

FP55

FP56

56.59 0.76 17.67 6.13 0.10 3.72 3.05 0.57 3.53 0.11 6.85 99.08

51.57 0.80 17.94 6.55 0.13 4.79 4.00 0.45 3.92 0.10 9.45 99.70

51.64 0.82 19.09 7.97 0.10 3.93 2.84 0.41 4.38 0.09 8.43 99.70

14 126 94 21 44 147 127 31 193 18 487 57 86 9

17 135 105 24 51 176 141 36 231 20 450 44 88 8

14 144 113 24 61 190 170 32 186 21 529 57 102 8

4.04 0.65 0.61 1.30 N.D N.D N.D

2.63 0.48 0.42 0.86 N.D N.D N.D

3.96 0.57 0.50 0.93 N.D N.D N.D

FP57 63.43 0.83 17.35 6.99 0.01 2.14 0.36 0.18 4.95 0.10 3.32 99.66

16 104 83 17 39 157 103 39 315 18 300 36 74 8

2.20 0.49 0.43 0.92 95.94 67.16 92.90

FP58 56.22 0.79 18.31 6.64 0.10 3.28 2.91 0.66 3.89 0.11 6.58 99.49

12 124 92 20 44 171 108 35 215 20 599 48 93 9

3.93 0.75 0.52 1.09 N.D N.D N.D

Cdo. Calderón

Sierra Espuña

FP59

FP68

59.10 0.97 19.37 6.38 0.02 2.45 0.25 0.33 5.83 0.16 5.02 99.88

16 125 103 13 40 180 418 45 337 22 541 65 106 7

3.96 0.43 0.63 1.63 96.05 66.12 92.79

FP60 61.13 0.93 17.70 7.37 0.02 2.25 0.24 0.28 5.41 0.16 3.65 99.14

18 124 96 13 38 157 384 43 368 22 497 55 104 7

2.99 0.36 0.57 1.45 96.13 65.84 92.83

52.90 0.78 18.68 7.62 0.06 2.71 3.27 0.17 5.17 0.05 8.14 99.55

15 132 103 21 38 196 85 27 154 16 884 54 86 6

3.60 0.37 0.52 1.42 95.53 75.56 93.93

FP69 51.68 0.80 20.84 7.07 0.02 3.06 2.14 0.19 6.26 0.12 7.26 99.44

20 146 113 18 41 238 111 30 138 19 894 45 92 8

2.25 0.40 0.40 1.10 93.94 73.26 91.55

FP70 51.14 0.80 19.03 7.96 0.06 2.77 3.92 0.16 5.30 0.05 8.54 99.73

17 130 102 22 42 199 100 29 146 17 2318 55 88 7

3.33 0.41 0.54 1.31 93.50 74.18 91.21

FP71

FP73

FP74

FP75

FP76

35.92 0.57 13.49 5.57 0.06 4.22 16.19 0.15 4.06 0.06 19.24 99.53

53.54 0.73 17.56 6.87 0.04 4.38 3.05 0.16 4.93 0.16 8.32 99.74

44.89 0.71 16.38 6.63 0.07 5.01 8.62 0.16 5.28 0.06 12.01 99.82

45.85 0.51 9.98 3.39 0.10 8.66 11.15 0.10 2.39 0.06 17.40 99.59

51.47 0.75 17.16 6.90 0.05 4.00 4.63 0.16 4.53 0.10 10.02 99.77

10 95 78 16 42 165 97 18 115 13 6087 40 54 6

15 112 87 20 40 186 135 39 223 19 437 49 79 6

14 110 91 18 38 201 89 27 170 17 403 43 83 6

10 62 45 12 30 102 81 31 220 13 705 30 63 5

15 122 91 18 42 186 108 31 182 20 448 47 82 6

4.00 0.61 0.51 0.95 N.D N.D N.D

3.27 0.39 0.56 1.23 N.D N.D N.D

3.03 0.44 0.47 1.13 N.D N.D N.D

3.00 0.53 0.67 1.00 N.D N.D N.D

3.11 0.39 0.52 1.12 N.D N.D N.D

Note: CIA, Chemical Index of Alteration (Nesbitt and Young, 1982); CIW, Chemical Index of Weathering (Harnois, 1988); PIA, Plagioclase Index of Alteration (Fedo et al., 1995). N.D. = not determined.

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Oxides (wt.%) SiO2 55.55 TiO2 0.98 Al2O3 20.86 Fe2O3 8.35 MnO 0.07 MgO 2.42 CaO 0.92 Na2O 0.25 K2O 4.51 P2O5 0.07 LOI 5.75 Total 99.73

Upper Triassic Arroyo de Zapateros

F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Fig. 10. (A) Major and (B) trace element distributions normalized to PAAS (Taylor and McLennan, 1985). The dark patterns are the range of samples.

Paleocene–Eocene package was estimated as about of 1418 m. Other packages made of harder lithologies (carbonates, conglomerates,…) suffered less compaction, as in the case of the Jurassic–Cretaceous, with a measured thickness of 350 m that becomes 635 m. Also the Oligocene sequence plus the Perona Unit are made of hard carbonates and conglomerates, and suffered a minor loading due to relatively soft sediments deposited during the early Miocene sequence; in this case the thickness only change from the 1250 m measured to 1418 m calculated in origin. 9. Discussion 9.1. Provenance The high compositional maturity of the sandstones, with dominant monocrystalline and polycrystalline quartz, scarcity of feldspar, and the presence of metamorphic and sedimentary lithic fragments and the mudstone compositional data from the continental redbeds

19

indicate that the Malaguide units had their main source in metamorphic and siliciclastic rocks, with a contribution of acid magmatic rocks. These sandstones plot mainly at the QmLt side in a QmFLt diagram and near Qt in the QtFL diagram (Fig. 7), reflecting their transition between a craton and a recycled orogenic provenance type (e.g. Dickinson, 1985). The basement of the Malaguide Complex mainly consists of a pre-Ordovician to Upper Carboniferous siliciclastic, siliceous and metasedimentary (slate, phyllite and quartzite) sequence including thin carbonate lenses and conglomerate bodies with some clasts of acid plutonic rocks (Vera, 2004). Equivalent successions are also found in other Alpine–Mediterranean terrains equivalent to the Malaguide Complex, such as the Moroccan Ghomarides, the Algerian Kabylides (Wildi, 1983), or the Italian Calabrides (Bonardi et al., 2001). In addition, in the Italian and Algerian counterparts of the Malaguide Complex, large late-Variscan plutonic bodies intrude the Palaeozoic sequence (Bonardi et al., 2001). Thus, source lithologies of the Malaguide Triassic were likely associated with this kind of terrains. The subordinate presence of felsitic volcanic detritus suggests an additional source, from Palaeozoic, probably Permian, felsic volcanic bodies. Also, in the sample set coming from the Upper Triassic beds of Sierra Espuña, an additional significant contribution from carbonate rocks is shown by the abundance of extrabasinal carbonate grains in sandstone and dolomitic pebbles in conglomerate bodies (Fig. 6C–D). Bulk rock and clay mineralogical data, as well as geochemical data, from the associated terrigenous mudstones also point towards a similar provenance, from continental crust source rocks of recycled orogenic or cratonic areas. The range of elemental ratios (La/Co, La/Cr, and La/Ni) of all samples studied fall in the range of similar published data (e.g. Cullers, 2000), suggesting a provenance from fairly silicic rather than basic source-areas. The high concentration of barium in some samples of the Sierra Espuña section indicates that hydrothermal and metasomatic processes could locally play some significant role, possibly related to partial replacement of some rocks as a result of interaction with hydrothermal solutions or other fluids (e.g. Hetherington et al., 2008 and references therein). In this sense, samples with high Ba values (Sierra Espuña) are close to evaporitic and carbonatic beds with some evidence of syn-sedimentary faulting, which has also been observed in the Malaga area (Cerrado de Calderón). In addition, basic volcanic and sub-volcanic rock bodies are associated with the Saladilla Fm redbeds in several areas of the Malaguide Complex (although they were not observed in any of the sections here studied). The Cr/V ratio in fine-grained sediments is an index of the enrichment of Cr over the other ferromagnesian trace elements, whereas Y/Ni monitors the general level of ferromagnesian trace elements (Ni) compared to a proxy for HREE (Y). Mafic–ultramafic sources tend to have high ferromagnesian abundances and such a provenance would result in a decrease in Y/Ni. In our case the Cr/V and Y/Ni proxies are close to the PAAS values, excluding a significant mafic–ultramafic supply (e.g., Mongelli et al., 2006) to the studied redbeds.

Table 5 Central ages calculated using dosimeter glass CN5 and ζ-CN5 = 366.5 ± 3.5. ρs: spontaneous track densities (×105 cm−2) measured in internal mineral surfaces; Ns: total number of spontaneous tracks; ρi and ρd: induced and dosimeter track densities (×106 cm−2) on external mica detectors (g = 0.5); Ni and Nd: total numbers of tracks; P(χ)2: probability of obtaining χ2-value for ν degrees of freedom (where ν = number of crystals − 1); a probability > 5% is indicative of an homogenous population. Samples with a probability b 5% have been analyzed with the binomial peak-fitting method. Sample number

FP50ar FP87ar

No. of crystals

11 20

Spontaneous

P(χ)2

Induced

ρs

Ns

ρi

Ni

0.27 1.37

8 78

0.34 1.47

103 837

92.4 51.3

Dosimeter

Age (Ma) ± 1 s

ρd

Nd

0.92 0.92

4368 4352

13.1 ± 4.8 15.6 ± 1.9

Note: Central ages calculated using dosimeter glass CN5 and ζ-CN5 = 366.5 ± 3.5. ρs: spontaneous track densities (×105 cm−2) measured in internal mineral surfaces; Ns: total number of spontaneous tracks; ρi and ρd: induced and dosimeter track densities (×106 cm−2) on external mica detectors (g = 0.5); Ni and Nd: total numbers of tracks; P(χ)2: probability of obtaining χ2-value for ν degrees of freedom (where ν = number of crystals − 1); a probability > 5% is indicative of an homogenous population. Samples with a probability b 5% have been analyzed with the binomial peak-fitting method.

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Table 6 Obtained thickness for each sediment package after successive steps of backstripping and decompaction, Sierra Espuña. Sequences

Measured thickness

Corrected thickness

Backstripping 1

Backstripping 2

Backstripping 3

Backstripping 4

Early Miocene Oligocene + Perona Unit Paleocene–Eocene Jurassic–Cretaceous Triassic Depth/subsidence

450 1250 275 350 700 3025

450 1418 1092 635 1775 5370

1418 363 387 801 4920

1092 485 1086 3502

635 m 1490 m 2410 m

1775 m 1775 m

m m m m m m

m m m m m m

9.2. Source-area weathering, and climatic features The chemical composition of clastic sediments depends on several factors including source-area composition, palaeoweathering, sorting,

m m m m m

m m m m

and, in some cases, burial history. Thus, in using terrigenous sediments to monitor provenance, one is faced with the nontrivial problem of evaluating and minimizing the effects of the many factors that can affect the chemical composition. The most widely used chemical index to

Fig. 11. (A) Subsidence analysis peformed for the Morrón de Totana Unit (Sierra Espuña area) taking into account the thicknesses of the Triassic to Miocene succession (more details in Appendix A). (B) Corrected subsidence using a standard backtripping method (for more details see Appendix A) by means of the program “Loadcap” by “Geostru Software” licensed to the Alicante University (reference no. G38RJ2). The uncorrected subsidence from Fig. 11A is also shown (red line).

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Fig. 12. Ternary (A) A–CN–K (Nesbitt and Young, 1982) and (B) A–C–N (Harnois, 1988) plots. Key: A, Al2O3; C, CaO; N, Na2O; K, K2O; Gr, granite; Ms, muscovite; Il, illite; Ka, kaolinite; Ch, chlorite; Gi, gibbsite; Sm, smectite; Pl, plagioclase; Ks, K-feldspar; Bi, biotite. The samples fall in the A–CN–K diagram close to the A–K join along a trend, indicating K addition during diagenesis.

determine the degree of source-area weathering is the Chemical Index of Alteration (CIA; Nesbitt and Young, 1982). In general, CIA values in Phanerozoic shales, ranging from 70 to 75, reflect muscovite, illite and smectite compositions, and indicate a moderately weathered source, whereas CIA values close to 100 are due to more intense weathering, which produces residual clays enriched in kaolinite and Fe–Al oxyhydroxides. The CIA values of analyzed redbeds are very low (average = 74 ± 1.2 for the eastern Betics; average = 68 ± 2.0 for the western Betics). In the A–CN–K triangular diagram the samples plot in a tight group on the A–K join, and close to the muscovite point (Fig. 12A). The sample distribution suggests a trend starting from a ‘granitoid source’ weathering with subsequent K-enrichment during diagenesis, in agreement with the mineralogical results (arrow of Fig. 12A). Since the CIA index is not sensitive to the weathering degree when K reintroduction occurs in the system, as in the present case, alternative indices can be used to monitor paleo-weathering at the source. Harnois (1988) proposed the CIW index (Chemical Index of Weathering), which is not sensitive to post-depositional K-enrichments. The redbeds show very uniform CIW values (average= 94± 1.1 for the eastern Betics; average = 94 ±3.3 for the western Betics) and these values, plotted in the A–C–N diagram (Fig. 12B), forming a tight array close to the A apex. This indicates intense weathering under steady-state conditions in the source area, where material removal rate matches the production of mineralogically uniform weathering products, generated in the upper zone of soil development. The degree of the chemical weathering can also be estimated using the Plagioclase Index of Alteration (PIA; Fedo et al., 1995). The studied redbeds show very high PIA values (average = 92 ± 1.5 for the eastern Betics; average = 91 ± 3.9 for the western Betics), indicating that most of the plagioclase has been converted to clay minerals. This, in turn, accords with the data obtained using the CIW index, and also indicates intense weathering at the source area. The A–CN–K triangle also can be used to constrain initial compositions of source rocks. The diagram illustrates metasomatized sample suite that typically has a linear trend with a less steep slope, but its intersection with the feldspar join (Fedo et al., 1995) indicates the likely source rock composition because the amount of K, Na and Ca addition to clays necessarily decreases as the amount of host clay material decreases. The studied samples seem to be related to granitoid source-rocks where the variable degree of K metasomatism in the samples reflects the quantity of secondary kaolin (derived from plagioclase) available for conversion to illite (Fedo et al., 1995).

9.3. Sorting and recycling It is well known that transport and deposition of clastic sediments involve mechanical sorting. Its effect on the chemical composition of terrigeneous sediments is important and may affect the distribution of paleoweathering and provenance proxies. The distribution of the chemical components within a suite is mainly determined by the mechanical properties of the host minerals. The process basically fractionates Al2O3 (clay minerals) from SiO2 (quartz and feldspars). Among the major and minor elements, the Al, Ti and Zr are generally considered the least mobile during chemical weathering. Significant amounts of Ti and Zr are typically fixed in resistate minerals such as zircon, rutile and ilmenite. Variations in these elements are conveniently expressed in a ternary Al–Ti–Zr plot (García et al., 1991). The interpretation of the compositional variation in this figure is based on the premise that sedimentation involves weathering, transport, mixing from different sources, and sorting (Sawyer, 1986). In the first three processes, the contents of insoluble elements, such as Al, Ti and Zr, may vary in response to the degree of leaching of the soluble elements. However, their relative proportions are transferred, from the source area, into the bulk sediment or regolith with little, or without, modification. This material is then sorted according to the hydraulic properties of its mineral components and it generates chemical fractionations between complementary shales and sandstones (García et al., 1994). In the studied case, a mixing trend is clearly envisaged (Fig. 13A), and it is mostly characterized by changes in the Al2O3/Zr ratio that could be due to a recycling effect (e.g., Caracciolo et al., 2011 and references therein). The presence of sorting-related fractionations is also evaluated when the Zr/Sc ratio (a useful index of sediment recycling; e.g. Cox et al., 1995; Hassan et al., 1999), is plotted against the Th/Sc ratio (indicator of chemical differentiation; McLennan et al., 1993). The mudrocks are not clustered along the primary compositional trend but fall along a trend involving zircon addition and thus sediment recycling (Fig. 13B). Recycling could significantly affect the weathering indices, which likely monitor a cumulative effect (e.g., Caracciolo et al., 2011 and references therein). This includes a first cycle of weathering at the source, which produced the siliciclastic Malaguide Palaeozoic sediments and equivalent rocks in other areas. The chemical weathering of such rocks, under hot, episodically humid climate with a prolonged dry season, would produce illitization of silicate minerals, oxidation of iron, and concentration of quartz in thick soil profiles that were later denudated by fluvial erosion, originating relatively mature,

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quartz-rich red deposits. The wet–humid conditions favored the formation of stream channels that eroded the soil profiles, whereas the dry season favored the sedimentation and pedogenic reddening (rubefaction). Successively, new wet–humid conditions cause the erosion of the sediments formed before. The alternation, in the Middle Triassic–Early Jurassic time interval, between these two different climatic conditions favored the recycling process affecting the studied sediments, as shown in the Al2O3–TiO2–Zr diagram and Zr/Sc vs. Th/Sc plot (Fig. 13A). 9.4. Malaguide Triassic palaeogeography, and evolution of the relief in the source area As it is well demonstrated for other Palaeozoic terrains of the Alpine Belt (von Raumer, 1998; von Raumer et al., 2003) recent studies in the Betic Cordillera have demonstrated that the Malaguide Palaeozoic succession was strongly involved in the Variscan Orogenesis (Martín-Algarra et al., 2009a, 2009b). Consequently, at the end of the Palaeozoic the Malaguide and equivalent successions in other Western Mediterranean Alpine Belts made part of a mountain belt that was raising and starting to be eroded in latest Carboniferous to earliest Permian time, and progressively denudated during the Permian and Triassic time. In the Malaguide area no Permian sediments have been yet dated with certainty and available biostratigraphic data only demonstrate that the continental redbed sedimentation was certainly active during the Middle Triassic, in relation to continental rift basins related to the late-orogenic extensional collapse of the previously formed Variscan Orogen and the evolving Pangea break-up (Martín-Algarra et al., 1995). The available paleocurrent data from the Saladilla Fm redbeds on both sides of the Gibraltar Strait indicate that alluvial sediments were deposited by flows coming predominantly from eastern and southern areas, accordingly to present-day coordinates: from the S and SE in the studied Malaguide outcrops in the Betic Cordillera (Roep, 1972; Mäkel, 1982; Martín-Algarra et al., 1995), and from the E and NE in the case of equivalent succession of the Moroccan Rif (Maate, 1996; Zaghloul et al., 2010). Even if strong vertical-axis palaeomagnetic rotations related to the Miocene formation of the Gibraltar arc are considered, rotated palaeocurrent data confirm eastern provenances for the Saladilla Fm redbeds, as palaeomagnetic data indicate clockwise and counterclockwise rotations in the Betic and Rifian outcrops, respectively (Platzmann, 1992; Osete et al., 2004). Consequently, the Permian– Triassic mountain area that acted as source for the Saladilla Fm redbeds

was located toward the E of the Malaguide Realm, closer to Neotethyan oceanic areas (Stampfli and Borel, 2004) than other regions of Iberia and NW Africa (Martín-Algarra et al., 1995). The rocks in such mountains were Palaeozoic metasediments deformed during the Variscan Orogeny, and associated with magmatic late Variscan and older intrusives, equivalent to those widely present in the Internal Domain units of the Algerian Kabylides (Wildi, 1983) and, especially, of the Italian Calabrides (Bonardi et al., 2001, and references therein). This palaeogeographic location of the Malaguide Realm (Fig. 14), nearby to regions where the Neotethys was rapidly spreading during the Triassic, helps to explain why evaporitic sediments are much less abundant and thinner in the Triassic successions of the Malaguide (and Alpujarride) Realm(s) than in the Triassic of the External Domain, where hectometer- to kilometer-thick evaporite succession are common, especially during the Late Triassic (Pérez-López and Pérez-Valera, 2007, and references therein). As the Malaguide Realm and associated reliefs were located close to Neotethyan ocean areas nearby the Triassic Equator (Stampfli and Borel, 2004), they could be more easily affected by trade winds that, coming from the ocean, would supply humidity to the mountains of the continent border, thus favoring wetter local climates and providing water supply needed for alluvial systems to flow from mountains to lowlands located towards the continent inside, as it was the case of the Malaguide Realm. As regards to the evolution of the relief in the source area, some information can be obtained from the sedimentological and sequence stratigraphical evolution. The sedimentary facies of the Triassic Saladilla Fm redbeds start with coarse-grained terrigenous sediments that become finer-grained upwards. This indicates the previous formation, during the Permian, of high relief source areas, which were actively weathered under warm climate in Permian–Triassic time, and rapidly denudated to form, first, alluvial fans that, later, evolved to fluvial systems with well defined channeled and flood plain depositional environments during the Middle Triassic. The facies evolution from alluvial to fluvial to shallow marine sedimentary environments during the Middle Triassic, and the generalization of coastal to shallow marine environments in most areas of the Malaguide Complex during the Late Triassic, also reflects the progressive peneplanation of the source area under warm and episodically humid climate. A similar evolution of the relief of the source area, as deduced from the sequence stratigraphical and facies evolution of the basin infill, can be also inferred from sandstone petrography and clay mineral evolution through time. The predominance of quartzose petrofacies in the

Fig. 13. (A) Ternary 15Al2O3–300TiO2–Zr plot after García et al. (1994), showing possible sorting effects. (B) Th/Sc vs. Zr/Sc plot (after McLennan et al., 1993). Samples depart from the compositional trend indicating zircon addition suggestive of a recycling effect.

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sandstones and of illitic and I–S mixed-layered clays within the terrigenous mudstones, together with the abundance of iron oxyhydroxides responsible of the red colors and with the other petrographical and geochemical features above discussed, indicate intense weathering in the source area under warm climate, with rapid denudation of the soil profiles that allows systematic preservation of small amounts of feldspars, both within sandstones and mudstones, especially during the Middle Triassic. The detrital signal within the Upper Triassic terrigenous mudstones is yet evidenced by the predominance of quartz and illite, by the persistence of feldspar and by the increased presence of terrigenous kaolinite and chlorite in the upper part of the studied succession, whereas the hot and episodically humid climate is also shown by the abundance of iron oxyhydroxides and the presence of kaolinite and increased amounts of I–S mixed layers. Under such climate conditions, the existence of a source area with low relief would allow stronger chemical weathering with intense leaching of silicates and progressive formation of kaolinite (which is absent in the lower part of the succession) in the soils of the well drained rised areas, whereas the decrease in the illite content and the increased proportion of smectite within I–S mixed layers would reveal intense pedogenesis in progressively wider topographically lowered areas with badly drained soils. The denudation of soils in such low relief areas allows to explain the increased proportion of I–S mixed layers, the decrease of detrital illite content, the presence of small amounts of kaolinite and chlorite, and the increased proportion of carbonates within the Upper Triassic sediments.

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9.5. Diagenesis, and burial/thermal history The sandstone petrography allows to characterize the diagenetic processes and the evolution history of the studied redbeds during burial. In particular, interstitial clays and authigenic minerals are important components of these sandstones. The clay minerals studied, in particular, are grouped into: a) depositional clays incorporated as mud intraclasts by reworking of overbank fines by fluvial processes; b) mechanically infiltrated clays occurring chiefly as coatings of tangentially accreted particles or as pore fills; and c) authigenic clays occurring as pore-fills and pore-lining kaolinite, illite and chlorite. Mechanical diagenetic infiltration of clays and precipitation of authigenic quartz, K-feldspar, albite and kaolinite contributed to porosity reduction. Extensive kaolinization and dissolution of detrital feldspar, mica and clay pseudomatrix occurred probably under an eo-diagenetic meteoric regime. The reaction in which smectite is a reactant and illite a product is recognized to occur over a predictable range of depth in mudstones, and it is here used as a comparative “geothermometer”, jointly with the illite crystallinity value (IC). In estimating temperatures from clay minerals evolution in the samples of the Malaguide Units, the Basin Maturity Chart proposed by Merriman and Frey (1999) was adopted, since it shows correlation of reaction progress in the smectite–I/S–illite series and IC with temperature, it refers to a large variety of geological settings and, therefore, is more representative of the evolution of such minerals through time. The studied samples show a content of illitic layers in the I–S mixed layers in a range of 75–90% (R ordering = 1–3; Reickeweite

Fig. 14. Middle (A) and Late (B) Triassic paleogeographic sketch maps of the Westernmost Tethys area, modified and redrawn from Vera (2004) and Critelli et al. (2008).

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number). The IC value, measured on both air dried and ethyleneglycol solvated oriented mounts of the samples, are in a range of 0.65–0.75° Δ2θ, with a IC mean value of 0.72° Δ2θ. These values are compatible with high diagenetic conditions (upper limit of late diagenesis; Merriman and Frey, 1999). By integrating data relative to the percentage of illitic layers in I/S mixed layers with illite crystallinity values, the temperature experienced by the studied samples can be estimated as in the range of 100–160 °C, following the Basin Maturity Chart. Similar results were obtained by previous clay mineralogy studies made both in the Malaga area and in Sierra Espuña (Ruiz-Cruz, 1992; Ruiz-Cruz and Puga, 1992; Ruiz-Cruz and Moreno-Real, 1993; Ruiz-Cruz and Andreo, 1996; Ruiz-Cruz and Rodríguez-Jiménez, 2002; Abad et al., 2003). So, starting from the temperature estimates by clay-mineral-based geothermometers, a diagenetic/tectonic evolution corresponding to about 4–6 km of lithostatic/tectonic loading can be hypothesized for the studied successions. These results are in agreement with the burial history deduced from the subsidence analysis performed for the Sierra Espuña succession (Fig. 11). Actually, it indicates that this Triassic sequence becomes highly compacted at the end of the Malaguide sedimentary evolution (Early Miocene), when the maximum depth of the Palaeozoic–Triassic boundary was located around 5370 m below the surface, and the thickness of the Triassic succession was reduced down to less than one half of its estimated original thickness they had immediately after its sedimentation, due to the sediment compaction. Such diagenetic evolution during burial also implies that the highest temperatures reached by the studied rocks should account during the Early Miocene, and that the later cooling and exhumation of the Malaguide Complex should be very rapid, as Middle Miocene marine sediments are widespread on folded Malaguide formations just N of the Sierra Espuña reliefs (Martín-Martín, 1996). This is consistent with thermochronological data about the late orogenic tectonometamorphic evolution of the Internal Domain all along the Betic Cordillera and, in particular, with that of the Alpujarride Complex (see Zeck et al., 1992; or Monié et al., 1994, among many others). The fission-track analysis to the two samples studied provides stronger constraints to the burial temperatures, and indicates that the temperatures reached were higher than total resetting temperature (about 130 °C). The calculated central age (15.6 Ma, early Langhian) should then reflect the time of cooling below the ca. 110 °C isotherm and the exhumation of Triassic deposits of the Malaguide Complex near to the surface, which is interpreted to have been very rapid. Although the possible existence of post-closure annealing cannot be totally excluded, as any statistically meaningful track length distribution study could be performed, the comparison with the clay mineralogical data suggests that no major burial affected the Malaguide samples here studied after their exhumation, however. Therefore, the obtained data can support that the closure temperature (ca. 110 °C) of the apatite system and the rapid cooling and exhumation of the studied rocks accounted during the Langhian, both in the western and eastern outcrop areas of the Malaguide Complex, in agreement with other apatite FT data obtained in deeper tectonic units below the Malaguide Complex (Esteban et al., 2005, among others). 10. Palaeogeographic summary and concluding remarks The Permian, Triassic and Early Jurassic time span is characterized by seasonal climate alternation that favored recycling processes. This period was also the time of rifting between Gondwana and Laurasia that finally opened the Tethyan ocean basins. The rifting started in the Permian, and progressed rapidly in the easternmost Tethyan areas with the rapid northward drift of the Cimmerian continent and with the active Triassic seafloor spreading that opened the Neotethys Ocean (Sengör, 1984; Cavazza et al., 2004). The rifting expanded

westward during the Triassic–Jurassic and generated continental rift basins. During the Jurassic it was followed by the detachment of microplates (e.g., the ‘Mesomediterranean Microplate’) from the main Europe and Africa plates, and by the opening of wide continental margins and narrow oceanic basins that characterize the Western Tethys region (Guerrera et al., 1993, 2005; Perrone et al., 2006; Critelli et al., 2008). The transition from the Late Triassic, characterized by a low firstorder sea-level stand and a time of high continental emergence, to the Early Jurassic, which was the time of the initiation of the firstorder sea-level rise in the mid-Mesozoic, reflects the evolution of the newly formed sedimentary basins (Golonka, 2000, 2002). The Early Jurassic was generally marked by a transgressive trend that would continue throughout the entire Jurassic. Consequently, the Malaguide Complex, which made part of the inner northern margin of the Mesomediterranean Microplate, was characterized by a Triassic– Jurassic stratigraphic evolution from continental–transitional redbeds to shallow marine, slope, and finally deep-marine carbonate/clastic sequences. Palaeoclimatically, the Permian to earliest Jurassic time was characterized by a transition from icehouse to greenhouse conditions (Frakes et al., 1992; Sellwood and Valdes, 2007; Stephenson et al., 2007; Twitchett, 2007). In subequatorial–tropical areas, like those corresponding to the palaeolatitudes of the future Mediterranean regions in general, and of the Malaguide Realm in particular, the continental rocks of the Mesomediterranean reliefs (formed after the Variscan Orogeny and during the starting rifting episodes related to the post-Variscan evolution) were subjected to chemical weathering under hot, humid climate with prolonged dry seasons. This produced illitization of silicate minerals, oxidation of iron and concentration of quartz in thick soil profiles that were later denudated by fluvial erosion, producing relatively mature, quartz-rich red deposits. The episodically wet-humid conditions favored the formation of stream channels that eroded the soil profiles, whereas the dry season favored the sedimentation, pedogenesis and rubefaction, and the successive wet–humid conditions caused the erosion of the sediments formed before (Mongelli et al., 2006). In the Triassic to Early Jurassic time, the alternation between these two different climatic conditions favored the development of the studied sediments, characterized by an important syn-sedimentary and early diagenetic evolution, as shown by mineralogy, chemistry and petrography of the studied Malaguide Triassic redbeds. As regards the late diagenetic evolution and the burial history, the clay-mineral features and the sandstone mineralogy of authigenic interstitial components evidence that redbeds experienced estimated temperature in the range of 100–160 °C, and a lithostatic/tectonic loading of about 4–6 km depending on the areas, which were both reached during the Early Miocene orogenic evolution. Consequently, the orogenic deformation of the Malaguide Complex occurred near to the surface, in upper structural levels. In a Western Mediterranean context, the Malaguide redbeds show age, lithology, sedimentary, diagenetic and orogenic evolution very similar to those characterizing the lowest Mesozoic sediments of several units, originated from Internal Domains of the Apenninic and Maghrebian Chains, at present cropping out in Kabylias, Calabria– Peloritani Arc and Tuscany (e.g., Perri et al., 2005a, 2005b, 2008a, 2008b, 2011; Somma et al., 2005; Mongelli et al., 2006; Perri, 2006, 2008; Perrone et al., 2006; Critelli et al., 2008; Zaghloul et al., 2010). Such close similarities reveal a common paleogeographic and paleotectonic evolution from the Middle Triassic up to the Miocene compressional deformation, and suggest a deposition in a same Mesozoic palaeogeographic realm, defined as Pseudoverrucano Sub-domain by Perrone et al. (2006). Finally, it is noteworthy that these rocks frequently have been compared with the Tuscan Verrucano (Cassinis et al., 1979), and frequently named as “Verrucano lithofacies”. This comparison

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cannot be totally shared because the Pseudoverrucano Sub-domain was contiguous to the Verrucano one, but the latter was more distal within the Mesozoic Mesomediterranean palaeomargin, and suffered a different Cenozoic tectono-metamorphic evolution (Perrone et al., 2006). Acknowledgments This research has been carried out within the MIUR-COFIN Project 2001.04.5835 “Age and characteristics of the Verrucano-type deposits from the Northern Apennines to the Betic Cordilleras: consequences for the palaeogeographic and structural evolution of the central-western Mediterranean Alpine Chains” (support to V. Perrone), MIUR-ex60% Projects (Palaeogeographic and Palaeotectonic Evolution of the CircumMediterranean Orogenic Belts, 2001–2005; and Relationships between Tectonic Accretion, Volcanism and Clastic Sedimentation within the Circum-Mediterranean Orogenic Belts, 2006; support to S. Critelli), and the 2006–2008 MIUR-PRIN Project 2006.04.8397 “The Cenozoic clastic sedimentation within the circum-Mediterranean orogenic belts: implications for palaeogeographic and palaeotectonic evolution” (support to S. Critelli). Financial support from the Projects CGL200909249 and CGL2011-30153-CO2-02 (MCI, Spain) and from the RNM208 and P08-RNM-03715 Research Groups of the Junta de Andalucia is also acknowledged. The authors are indebted to Robert L. Cullers, José Arribas and the Editor Andrew D. Miall for their reviews and suggestions on an earlier version of the manuscript. This article is dedicated to our dear friend Glauco Bonardi. Appendix A. Subsidence analysis and compaction procedures Subsidence analysis uses the backtripping method to quantify the tectonic versus thermal components of the subsidence of sedimentary basins during long time periods, by assuming a response of the crust based on local isostasy, and by taking into account the water load (if the basin is marine), the sediment load, and the irreversible loss of porosity of the sediments due to compaction during burial at increasing depth (Watts and Ryan, 1976; Steckler and Watts, 1978; Sclater and Christie, 1980). A plot is so performed to delineate the vertical movements of a reference stratigraphic horizon, and to track the subsidence/uplift history of the basin due to sediment load, tectonic load and sea-level change (Xie and Heller, 2009). The study has been performed using the program “Loadcap” by “Geostru Software” licensed to the Alicante University (reference no. G38RJ2). For that purpose, the program requires the thickness of each sedimentary sequence (isochronous packages), the mean saturated density of the considered lithotypes, and the mean edometric module, a parameter related to the stretching and the % of pores of the sediments or rocks and, hence, to their capability to be compacted. The program allows to insert the load due to deposition of each sedimentary sequence of the basin infilling, the stacked tectonic units and also the load due to the mean water column, calculated from paleobathymetric interpretation. The analysis must be carried out into three phases. First, the subsidence (due to flexure or faulting of the basement) is processed by analyzing the thickness of each sedimentary package. After, a backstripping analysis and decompaction of the sedimentary sequences (as they become increasingly compacted with time due to the sediment load of the successive layers), is done by using the Loadcap program. The backstripping is performed by isolating the isochronous packages one-by-one, being “peeled off”, or backstripped, and the lower bounding surface moved upward to a datum. By successively backstripping isochrons, the deepening history of the basin can be plotted in several steps, one for each “stripped off” sedimentary sequence. Finally, a correction by compaction of the sedimentary sequences allows to obtain the original thickness and subsidence of the sedimentary record of the study area.

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The parameters used to run the program for the Sierra Espuña area case study are enclosed in the table below. In the studied case, the considered sedimentary packages are: 1) Triassic; 2) Jurassic–Cretaceous; 3) Paleocene–Eocene; 4) Oligocene (including the tectonically stacked Perona Unit, which thrusted at that time onto the Morrón de Totana Unit: see Martín-Martín et al., 1997c; Martín-Martín and Martín-Algarra, 2002 for details); and 5) early Miocene (from the middle Miocene onwards the unroofing of the Sierra Espuña succession progressed rapidly due to tectonic folding and rising). The backstripping was performed by isolating the packages mentioned before one-by-one in four steps, one for each sedimentary sequence “stripped off”, starting from the youngest package; early Miocene, Oligocene–Perona Unit, Paleocene–Eocene and Jurassic–Cretaceous. The analysis uses decompaction data of the remaining sequence following each stage of the backstripping. This takes into account the original thickness before the loading and the maximum depth reached for each sedimentary sequence due to the loading of the following packages (Table 6 of the paper). To run the program, the mean density and the thickness of the loading sequence are required. The mean density used has been saturated since sediments take place in water realm, and taken from the above mentioned literature. For each deposited sequence, the thickness and the mean edometric module (not the density) are required. For successively older packages, the mean edometric modules have been calculated in two conditions: when the sequence is the last deposited it is considered not compacted, and when the sequences had been deposited before and have suffered compaction, diagenesis and lithification, they are considered as compacted. The mean density and the mean edometric module have been calculated according to the thickness of each layer composing the sedimentary sequence. As the thrusting of the Perona Unit started during the Oligocene and followed during the earliest Miocene, the effect of the tectonic stacking of this unit has also been calculated. The present-day substratum of the Morrón de Totana Unit is made of a thin, tectonically truncated, Palaeozoic sole under the Triassic and younger succession, but the Paleozoic sole of all Malaguide successions was much thicker before the Oligo-Miocene thrusting and nappe stacking. Consequently, and taking into account the maximum thickness of the pre-Mesozoic rocks of the Malaguide Complex observed near Malaga, where the whole Palaeozoic succession is preserved, a 2000 m thick rigid basement has been used for the subsidence analysis of the studied succession. Nevertheless, the thickness of substratum of the Morrón de Totana Unit (and also that of the underlying Malaguide and Alpujarride units) was (were) seriously modified since thrusting started. Consequently, and to correct this, during the Oligocene–Miocene the substratum of the analyzed succession has been considered to be formed, from bottom to top, by: a 2000 m thick rigid Palaeozoic basement, followed by a 1 km thick nappe stack made by the Triassic successions belonging to the underlying Alpujarride (Molinos) and Malaguide (Jaboneros, Yechar and La Santa) units followed by a 50 m thick Palaeozoic sole underlying the Meso-Cenozoic sediments. The edometric module and density used for the Palaeozoic rocks reproduce rigid conditions (Table 6 of the paper) whereas those used for the underlying Triassic successions of the nappe stack are the same parameters used for the Morrón de Totana succession. For all the sedimentary sequences studied in Sierra Espuña, a lateral continuity of 10 km has been considered for the packages studied. Finally, the hydrostatic load of the water column has been taken into account only in the case of deep marine sediments (Miocene turbidites related to the foredeep stage of the basin evolution: see Martín-Martín and Martín-Algarra, 2002, for details). This case only occurs during the early Miocene, when 500 m of water column has been considered according to benthic foraminifera-based paleobathymetric studies in equivalent deposits in nearby Malaguide areas. In the case of continental, transitional or shallow marine sediments (from the Mesozoic to the Paleogene) the loading of the water column is negligible.

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F. Perri et al. / Earth-Science Reviews 117 (2013) 1–28

Table: Parameters used for subsidence analysis of the Morrón de Totana succession (Sierra Espuña). Age sequences

Early Miocene

Perona Unit + Oligocene

Paleocene–Eocene

Jurassic–Cretaceous

Triassic

Palaeozoic

Rocks and sediments

Water column (500 m) Marls & silexites Conglomerates Silts & sandstones Carbonates (Perona Unit) Marls Conglomerates & marls Marls Clays Calcarenites Marly limestones & sands Limestones Dolostones Clays with gypsum Carbonates Clays, sands & sandstones Conglomerates & sandstones Slates, quartzites & greywakes

Saturated density (T/m3) 1 2.5 2.3 2 3.1 2.5 2.6 2.5 1.5 2.8 2.5 3 3.2 1.8 3 2.2 2.3 3

Edometric module (kg/cm2) Not compacted

Compacted

Only load, not compacted

100 1500 100 100 1000 750 1500 1500 80 1500 400 1350

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2000 120 1750 120 120 1500 850 2000 2000 100 1900 450 1500 2500

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