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Accueil Tous les numéros Volume 196 (2025) BSGF - Earth Sci. Bull., 196 (2025) 15 Full HTML
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Numéro
BSGF - Earth Sci. Bull.
Volume 196, 2025
Numéro d'article 15
Nombre de pages 26
DOI https://doi.org/10.1051/bsgf/2025008
Publié en ligne 4 septembre 2025

© H.R. Campos-Rodríguez et al., Published by EDP Sciences 2025

Licence Creative CommonsThis is an Open Access article distributed under the terms of the Creative Commons Attribution License (https://creativecommons.org/licenses/by/4.0), which permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly cited.

1 Introduction

Petrological, geochemical and geochronological studies of mafic rocks are crucial for identifying mantle sources, crustal contamination, and tectonic settings that prevailed in magmatic provinces (Chauvel et al., 2008; Zhao et al., 2009; Tappe et al., 2013; Orejana et al., 2020; Villaseca et al., 2022). In equivalent domains of the Variscan Belt (e.g. Armorican and Iberian Massifs, Fig. 1A), several studies on mafic suites emplaced prior and during the Variscan orogeny suggest the existence of different sources and an important interaction between mantellic and crustal magmas as recorded by their geochemical signatures (e.g. Dostal et al., 2019; Martínez Catalán et al., 2021; Schulmann et al., 2022; Villaseca et al., 2022). In the Armorican Massif (Fig. 1A), geochronological and geochemical studies on mafic dykes propose the occurrence of a magmatic event that took place during the early stages of the Variscan orogeny (∼360 Ma; Pochon et al. 2016b). The emplacement of these mafic units was related to a slab roll-back in a in a convergent back-arc geodynamic setting whose source was associated with the melting of an enriched lithospheric mantle (Pochon et al. 2016b; Pochon, 2018). In the case of the Central Iberian Zone (CIZ) of the Iberian Massif (Fig 1B), the emplacement of several mafic units has been linked to an extensional and convergent geodynamic setting associated with the opening and subduction of the Rheic ocean as well as the Variscan collision (e.g. Simancas et al., 2003, 2006; Gutiérrez-Alonso et al., 2008; Solá et al., 2010; Villaseca et al., 2022; among others). All of the sources of these mafic rocks have significantly evolved through time, involving a metasomatized mantle, crustal contamination and other relevant processes (e.g. Villaseca et al., 2022 and references therein).

Previous studies identified several magmatic events spanning from the Neoproterozoic (573 Ma, late Pan-African (Cadomian) cycle; e.g. Bandrés et al., 2004) to the early Permian (274 Ma, at the end of the Variscan Orogeny; e.g. Orejana et al., 2020), in various domains of the CIZ (Fig. 1B), such as: the Mérida Montoro Complex (Bandrés et al., 2004); the Spanish Central System (SCS, Barbero and Villaseca, 2000; Orejana et al., 2017, 2020; Villaseca et al., 2002, 2015); the Carrascal Massif (Solá et al., 2003; 2010); the basalts from El Castillo in the Valongo-Tamames Syncline (Gutiérrez-Alonso et al., 2008); the CS mafic rocks (Garcia de Figuerola et al., 1974; Carnicero and Castro, 1983; López-Moro et al., 2005, 2007, 2020); the AS mafic rocks (Saupé, 1990; Saupé and Arnold, 1992; Higueras, 1995; Higueras et al., 1995, 2013; Hall et al., 1997); the Avila Batholith in the SCS (Bea et al., 1999; Montero et al., 2004a; Scarrow et al., 2009); the Sierra Bermeja pluton (Errandonea-Martin et al., 2019); the Calzadilla pluton (López-Moro et al., 2017) and the Toledo Anatetic Complex (Bea et al., 2006). These mafic magmatic events in the CIZ are characterized by various geochemical features, such as the occurrence of enriched and metasomatized mantle sources (e.g. López-Moro et al., 2007; Orejana et al., 2009, 2017; Villaseca et al., 2015); as well as crustal recycling or crustal contamination (e.g. Montero et al., 2004; Orejana et al., 2009; de Almeida Coelho, 2021).

Despite the extensive work related to the mafic magmatism in the CIZ, the geochemical characteristics and emplacement ages of some of these mafic suites are still not totally constrained. This is the case for the mafic intrusions emplaced within the following three folded Palaeozoic metasedimentary sequences (Fig. 1B): The Codosera (CS), Almadén (AS) and Guadalmez synclines (GS). The CS hosts the biggest swarm of mafic sills reported in the Iberian Massif, which is spatially associated to one of the largest Sb-W districts in Spain (San Antonio Mine; e.g. Gumiel and Arribas, 1987). Mineralogical, structural and geochemical studies performed in this area propose that the mafic rocks were emplaced during an extensive event during late Devonian (Garcia de Figuerola et al., 1974; Carnicero and Castro, 1983; López-Moro et al., 2007, 2020). The mafic magma is suggested to derive from a metasomatized mantle presenting a significant crustal contamination (López-Moro et al., 2007, 2020). The AS hosts several mafic units and the world class Almadén Hg deposit. Based on geochronological data and field observations, the mafic magmatism in this structure is believed to have been mainly emplaced at early Silurian times followed by minor magmatic events from middle Silurian and lower Devonian (Higueras, 1995; Hall et al., 1997; Higueras et al., 2013; Palero-Fernández et al., 2015). Finally, the mafic rocks from the GS have not been studied yet.

In this study, we present new petrological and geochemical data for some of the mafic intrusions located in the CS, AS and GS. We constrain the emplacement ages of these mafic magmas by U-Pb dating on apatite. We analyze the geochemical evolution and the possible mantle sources of the Ordovician to Carboniferous mafic magmatism in the CIZ. In addition, we discuss the timing of the magmatism in the CS, AS and GS and then at the scale of the Variscan orogeny by comparing with equivalent domains in the Armorican Massif.

thumbnail Fig. 1

A) Geological sketch of the Variscan Belt showing the main terranes and domains, modified after Martínez Catalán et al. (2014); B) Simplified geological map of the Iberian Massif (Julivert et al., 1974; Simancas, 2019) showing mafic rocks in the studied areas and the mafic suites reported in the Central Iberian Zone (CIZ) from Ordovician times to the late stages of the Variscan orogeny. Abbreviations: OMZ: Ossa Morena Zone; GTMZ: Galicia-Tras-Os Montes-Zone; WALZ: West Asturian Leonese Zone (WALZ); CZ: Cantabrian Zone; MN: Montagne Noire; ECM: External Crystalline Massifs of the Alps; VM: Vosges Massif; BF: Black Forest; CP: Calzadilla Plutón (López-Moro et al., 2017); LA: Praia de Labruge (de Almeida Coelho, 2021); GD: Godomar (Medeiros et al., 1981); F: Farminhao (Cotrim et al., 2021); AP: Appinites (Montero et al., 2004b, a; Scarrow et al., 2009); TAC: Toledo Anatetic Complex (Bea et al., 2006); C: Caloco Sector (Villaseca et al., 2015; Orejana et al., 2017); R: Revenga Sector (Villaseca et al., 2015; Orejana et al., 2017); TM: Tenzuela Massif; MMC: Mérida Montoro Complex(Bandrés et al., 2004); SB: Sierra Bermeja (Errandonea-Martin et al., 2019); CM: Carrascal Massif (Solá et al., 2010); MG : Microgabbros (Orejana et al., 2020); CS: Codosera Syncline (López-Moro et al., 2007); AS: Almadén Syncline (Higueras, 1995; Higueras et al., 2013); GS: Guadalmez Syncline (Gumiel and Arribas, 1987; Lorenzo Álvarez et al., 1995).

2 Geological setting

2.1 The mafic magmatism in the Central Iberian Zone

The Central Iberian Zone (CIZ) belongs to the Iberian Massif (Fig. 1B) and is limited to the South by a NW-SE sinistral transpressive structure called the Tomar-Badajoz-Córdoba-Shear-Zone, which marks the boundary with the Ossa Morena Zone (Sanderson et al., 1991). It was overlapped at the north by thrust sheets of both oceanic and continental crust, belonging to the Galicia-Tras-Os-Montes Zone (Díez Fernández and Arenas, 2015; Arenas et al., 2016). The CIZ presents metasedimentary rocks that have experienced metamorphism of varying grade (from prehnite-pumpellyte to granulite and eclogite facies, e.g. Higueras, 1995; Barbero and Villaseca, 2000; Martínez Catalán et al., 2014; Azor et al., 2019). It includes a siliciclastic succession of Neoproterozoic to Cambrian age (Schist Greywacke Complex, SGC; Lotze, 1945), unconformably overlain by a Paleozoic sequence (composed by sandstones, shales, slates, minor limestones and quartzites). Within the early Paleozoic succession, outcrops a unit known as the “Ollo de Sapo Formation” of Cambrian-Ordovician age (495–470 Ma; García-Arias et al., 2018), composed of a variety of metaigneous and metavolcanics rocks whose origin has been related to a possible partial melting of the Schist Greywacke Complex (García-Arias et al., 2018).

During the Precambrian to Carboniferous period the metasedimentary sequences of the CIZ were affected by at least six mafic magmatic events, including (e.g. Villaseca et al., 2022):

  1. The late Neoproterozoic orogeny (Cadomian), which is characterized by the intrusion of rocks of calc-alkaline affinity from the Mérida Massif, dated at 557 ± 3.7 Ma (U-Pb ID-TIMS in zircon; Bandrés et al., 2004).

  2. The Ordovician period is marked by a rifting episode that led to the opening of the Rheic ocean (Matte, 2001; Arenas et al., 2007; Martínez Catalan et al., 2007; Nance et al., 2012). Mafic rocks emplaced during this period comprise: i) alkaline gabbros from the Carrascal Massif (471 Ma; zircon 207Pb/206Pb evaporation; Solá et al., 2010); ii) mafic rocks of tholeiitic affinity from the Spanish Central System in the Revenga- Caloco sectors and the Tenzuela Massif, (emplaced between 473 and 453 Ma; SHRIMP U-Pb zircon; Orejana et al., 2017; Villaseca et al., 2015); iii) metagabbros, amphibolites, and ultramafic rock of calc-alkaline affinity in Farminhao, northern Portugal (ca. 480 Ma; Cotrim et al., 2021).

  3. During the early Silurian period, an alkaline intraplate mafic event emplaced microgabbros, dolerites, porphyritic lavas and diatremes of basaltic composition with ultramafic xenoliths in the Almadén area and caused the hydrothermal activity that lead to an important Hg mineralizing event, responsible for the formation of the world class Almadén Hg deposit (Higueras, 1995; Higueras et al., 1995, 2013).

  4. From the Middle to Upper Devonian other alkaline mafic suites were intruded in the AS (e.g. Hall et al., 1997; Higueras et al., 2013; Palero-Fernández et al., 2015) and in the Valongo-Tamames syncline ( U-Pb zircon age of 394.7 ± 1.4 Ma; Gutiérrez-Alonso et al., 2008).

  5. The Upper Devonian to early Carboniferous period marked the closure of the Rheic Ocean and the beginning of the Variscan orogeny (Arenas et al., 2007; Martínez Catalan et al., 2007; Nance et al., 2012; Paquette et al., 2017). The magmatic events during this period lead to the emplacement of continental tholeiitic dolerites from La Codosera syncline (Carnicero and Castro, 1983; López-Moro et al., 2005, 2007, 2020), and dolerites from the Gondomar area in Portugal (Medeiros et al., 1981).

  6. During the last stages of the Variscan orogeny, in the late Carboniferous, the Variscan belt underwent extensional collapse in response to an adiabatic melting of the lithospheric mantle (e.g. Montero et al., 2004a, b), resulting in the intrusion of mafic and ultramafic rocks represented by: i) appinites, of the Avila Batholith (319 ± 3; U-Pb zircon; Montero et al., 2004a, b; 310 ± 3Ma, U-Pb zircon; Scarrow et al., 2009); ii) ultramafic rocks related to a mixing between mafic, ultramafic magmas and crustal material (Montero et al., 2004a, 2024b; Scarrow et al., 2009); iii) gabbros of the Toledo Anatetic Complex (dated at 307 ± 2 Ma by ion microprobe U-Pb on zircon; Bea et al., 2006); iv) Dyke swarms of microgabbros of calc-alkaline affinity (MG, Fig. 1B) intrusive in the eastern part of the Avila Batholith (299-292 Ma; SHRIMP U-Pb zircon; Orejana et al., 2020); v) dolerites from the Praia de Labruge in Portugal, West of Porto (299.8 ± 4.5 Ma; 40Ar/39Ar in mica; de Almeida Coelho, 2021).

2.2 La Codosera (CS), Almadén (AS) and Guadalmez (GS) synclines

2.2.1 La Codosera syncline (CS)

The CS crops out between the Badajoz and Cáceres provinces, in the center of the Extremadura region in Spain (Figs. 1B and 2). This NW-SE trending structure extends over a length of 100 km and a width of 10 km. This syncline was the result of the NE-SW shortening associated with the collision of Gondwana and Laurussia, which was responsible for the formation of the Variscan orogen between the Upper Devonian and lower Carboniferous (e.g. Martínez Catalan et al., 2007; Nance et al., 2012; Paquette et al., 2017). The main deformation documented in the studied area was dated in the Cabo Ortegal Complex of the CIZ at ca. 359 Ma (40Ar-39Ar on muscovite on the axial planar foliation; Dallmeyer et al., 1997) and is represented by a slaty cleavage (S1) and a down dip lineation indicating that the strike slip elements were not relevant and that the main deformation was perpendicular to the orogen (Martínez Catalán et al., 2009, 2014).

The CS is composed of a Palaeozoic sequence including Ordovician to Devonian quartzites, sandstones, schists and slates, overlain by a Carboniferous sequence of slates/shales, with minor sandstones and some calcareous levels; known as the Gévora Formation (Gumiel et al., 1976; Gonçalves and Perdigao, 1978; Santos Garcia and Casas Ruiz, 1979, 1982; Gumiel, 1982; Soldevila Bartolí, 1992; Rodríguez González et al., 2007). All of these successions unconformably overlie the so-called Schists and Greywacke Complex of late Precambrian age (San José et al., 1990).

The Gévora formation, located at the core of the CS (Fig. 1B), is intruded by several sill-like mafic bodies (dolerites, Fig. 2B) showing a N80°E-N120°E strike and a steep dip (López-Moro et al., 2005, 2007, 2020). Their thickness is very variable and range from a few meters up to 400 m, with a maximum length of 300 m (López-Moro et al., 2005, 2007, 2020). The intrusions have not been dated yet but are believed to have been emplaced during late Devonian to early Carboniferous times, based on stratigraphic observations (López-Moro et al., 2020).

The whole CS and all the mafic sill swarms are crosscut by the Messejana-Plasencia crustal fault (Figs. 1B and 2B), which was subsequently injected by the Messejana-Plasencia dolerite dyke, Upper Triassic in age (Rhaetian: 203 ± 2 Ma, 40Ar/39Ar on biotite; Dunn et al., 1998). This dyke is related to the beginning of the rifting stage that led to the opening of the central Atlantic Ocean (De Vicente et al., 2021). Subsequently, the Tethyan orogeny induced a N-S compression that accommodated a 5 to 20 km thick left-lateral strike-slip deformation belt along the Jurassic Messejana-Plasencia dyke, with the formation of strike-slip (transpressional to transtensionnal) and thrust basins (De Vicente et al., 2011, 2021). The Badajoz Cenozoic basin widens to the south of the CS and is a thrust or open-ramp basin (De Vicente et al., 2011) limited to the west by the Messejana-Plasencia fault and to the north the E-W North Badajoz thrust fault that crosscut the eastern part of the CS as well as its SE border (Figs. 1 and 2). Thus, the North Badajoz thrust fault is responsible for the uplift of the eastern part of the CS relatively to its western part. Therefore, deepest structural levels crop out to the East of the Messejana-Plasencia fault, whereas the shallowest structural levels crop out to the west of the fault. The effects of these structural settings within the CS are represented by: (i) A greater number of thicker doleritic sills found at the eastern part of the structure when compared to the western side and ii) the occurrence of numerous gold (Roberts et al., 1991; Dee and Roberts, 1993; Murphy and Roberts, 1997) and antimony hydrothermal systems, among them the San Antonio mine operated until 1986 (Gumiel et al., 1976; Arribas and Gumiel, 1984; Gumiel and Arribas, 1987), in the western part of the CS when compared to the eastern part.

thumbnail Fig. 2

A) Simplified geological map of the Iberian Massif showing the studied area of La Codosera syncline (CS). Modified after López-Moro et al. (2020) and references therein ; B) Geological map of the CS showing the mafic samples studied in this work. This map was created based on the Geological Maps (scale 1:50000) of Alburquerque (Santos García and Casas Ruiz, 1982), Puebla de Obando (López Díaz et al., 1989), Villar del Rey (Santos García et al., 1990) and Botoa (Insua Marquez et al., 1990).

2.2.2 The Almadén (AS) and Guadalmez (GS) synclines

The AS and GS (Fig. 3B) are located in the Ciudad Real and Cordoba provinces at around 260 km to the NE of Sevilla. Both represent WNW-ESE trending structures showing sub-vertical axial planes with lengths of 25–30 km long and widths of 12–15 km. The AS hosts the world class Almadén Hg deposit (Higueras et al., 1995, 2013; Hall et al., 1997; Hernández et al., 1999; Palero and Lorenzo, 2009; Palero-Fernández et al., 2015), whereas the GS hosts several Sb occurrences and mineralization (Gumiel, 1982; Gumiel and Arribas, 1987; Lorenzo Álvarez et al., 1995; Barquero et al., 2022). The AS and GS include a Palaeozoic sequence overlying the Precambrian succession known as the Schist-Graywacke Complex (San José et al., 1990). Like the CS, the whole sequence has been folded and affected by the ca. 359 Ma Variscan deformation phase responsible for a slaty cleavage in the CIZ (Higueras, 1995; Dallmeyer et al., 1997; Higueras and Oyarzun, 2005; Palero and Lorenzo, 2009; Higueras et al., 2013; Palero-Fernández et al., 2015). A very low grade metamorphism ranging from zeolite to prehnite-pumpellyite facies has been reported in the AS (Saupe, 1973; García San Segundo et al., 1985; Higueras et al., 1995), whereas metamorphism up to greenschist facies is reached in the GS (i.e. mineral associations ranging from quartz+muscovite to quartz+muscovite+chlorite+albite; Lorenzo Álvarez et al., 1995). In both synclines, the low grade metamorphism has been associated with the Variscan deformation (García San Segundo et al., 1985; Higueras et al., 1995; Lorenzo Álvarez et al., 1995).

Differently from the CS, quartzite is the most representative lithology in the Palaeozoic sequence of the AS and GS, including four main units dating back from the Lower Ordovician to the Lower Devonian (Higueras et al., 2013; Palero-Fernández et al., 2015). This lithology is accompanied by siltstones and slates, with intercalations of mafic volcanic and subvolcanic rocks (Almela et al., 1962; García San Segundo et al., 1985; San José et al., 1990; Lorenzo Álvarez et al., 1995).

The Carboniferous rocks in the Paleozoic sequence of the AS are essentially tillite, basalt and tuff with trachytic and rhyolitic compositions (García San Segundo et al., 1985). In the case of the GS, the Carboniferous sequence is represented slates and limestones (Pardo and García-Alcarde, 1984; Lorenzo Álvarez et al., 1995).

In the AS mafic units are mostly intrusive in the lower Silurian to Devonian rocks under the form of sills, dykes, volcano sedimentary rocks, porphyritic lavas and diatremes (breccia-tuffs) (García San Segundo et al., 1985; Higueras, 1995; Hall et al., 1997; Higueras et al., 2000, 2013; Palero and Lorenzo, 2009; Palero-Fernández et al., 2015). In the GS, the mafic rocks, mostly intrusive in Silurian shales, are present as small diabase sills (Lorenzo Álvarez et al., 1995; Barquero et al., 2022).

It has been proposed that the mafic volcanism in the AS acted as a trigger for the hydrothermal activity resulting in the formation of the world class Hg deposits of the Almadén district (Higueras, 1995; Higueras et al., 2000, 2013; Higueras and Oyarzun, 2005; Palero-Fernández et al., 2015). Hall et al. (1997) performed 40Ar/39Ar dating on illites associated with the last Hg mineralization stages and Cr-rich micas related to earlier mineralization stages, yielding ages at ca. 360 Ma and 427-365 Ma, respectively.

thumbnail Fig. 3

A) Simplified geological map of the Iberian Massif showing the studied area of the Almadén (AS) and Guadalmez (GS) synclines. Modified after López-Moro et al. (2020) and references therein ; B) Geological map of the of the Almadén (AS) and Guadalmez (GS) synclines showing the mafic samples studied in this work. This map was created based on the Geological Maps (scale 1:50000) of Chillón (Lorenzo Álvarez et al., 1995), Almadén (García San Segundo et al., 1985), Hinojosa del Duque (Mira López et al., 1983a) and San Benito (Mira López et al., 1983b).

3 Analytical methods

3.1 Sampling

Seventeen dolerite intrusions were sampled in the three studied synclines (Figs. 2 and 3). Ten samples from the CS were located within the lower Carboniferous Gévora Formation (Rodríguez González et al., 2007), (Fig. 2; SA2008, SA2013, SA2014A, SA2014B, SA2015, SA2022A, SA2022B, SA2022C, SA2026 and SA2042). Four samples from the AS, were collected from the volcano-sedimentary rocks of Upper Devonian age, composed of tuffs and volcanic breccias (Fig. 3; AL2076, AL2077A, AL2077B and AL2088), and three samples from the GS were collected from an Upper Ordovician to lower Silurian sedimentary sequence, represented by the units known as Pizarras de Muro, Criadero Quartzite and ampelitic slates (Fig. 3; AL2072, AL2073, AL2075).

3.2 Whole rock geochemistry analyses and mineral chemistry

Whole rock geochemistry analyses of the 17 mafic rocks were performed at ALS laboratories in Dublin (Ireland), employed methods are described in the Supplementary material 1. The whole rock geochemistry analyses with the detection limits for the samples in this study can be found in Supplementary material 1 and 2. The major element chemistry of the dated apatite crystals was analysed with a Cameca SXFive electron probe micro analyzer (EPMA) at the BRGM/CNRS/ISTO facilities, France. Method description can be found in Supplementary material 1.

3.3 U-Pb dating on apatite

Apatite is the most common U-bearing mineral in mafic rocks and may represent an important asset for U-Pb dating of small mafic bodies like sills or dykes, which experienced a fast cooling and were not heated above the closure temperature of apatite (375–550 °C, Cochrane et al., 2014; Schoene and Bowring, 2007). As apatite can incorporate non-negligeable amounts of common Pb in its crystal structure, and because it usually has low U and therefore low radiogenic Pb contents, the resulting radiogenic Pb/common Pb ratios are usually low (Hughes and Rakovan, 2015).

The geochronological analyses were performed in both thick-polished sections of 150 μm and separated apatite crystals from seven dolerite samples. Selected apatite grains were imaged by cathodoluminescence (CL) at the ISTO-BRGM facilities, in order to observe possible zoning and identify the best locations for LA-ICP-MS analyses. We used a Cathodyne-OPEA cold cathode instrument, in which the samples are placed into an argon pressure of ∼50 mTorr, and irradiated by an electron beam with ∼15 kV and ∼100 mA. The angle between the axis of cannon and the horizontal is 25°.

We performed U-Pb geochronology in the LA-ICP-MS facilities of the GeOHeLiS platform at the University of Rennes. The LA-ICP-MS system is equipped with an Excimer ESI NWR193UC coupled to a Q-ICP-MS Agilent 7700×.

Instrument calibration was carried out before each batch of analyses by monitoring the signal of the 238U, the ratio between the uranium and thorium and by minimizing the ThO+/Th+ ratio (<0.5%) using a NIST SRM 612 reference glass. All the material produced by the ablation is carried by He then mixed with Ar and N2 before being injected into the plasma source. The signals of 206Pb, 207Pb, 208Pb, 238U, 204(Pb+Hg) and 43Ca are collected during each analysis. The 235U is calculated based on the 238U/235U ratio of 137.818 (Hiess et al., 2012).

Each measurement consisted of 20 seconds of background collection, followed by 60 seconds of ablation, and 15 seconds of wash-out delay. The laser pulse rate was 5 Hz. In the first session a rectangular spot was used (30 μm × 40 μm) for samples SA2022C, SA2026 and AL2073. In the second session, two types of spots were used, the first one consisted of a rectangular spot of 25 μm × 50 μm for samples AL2072 and AL2075; whereas the second one was a round spot of 40 μm diameter used in samples SA2014A and AL2076.

The sequence of measurements was the following: two measurements of the Madagascar apatite standard (ID-TIMS age of 473 ± 0.7 Ma, Thomson et al., 2012), one measurement of the Durango apatite standard (31.44 ± 0.18 Ma, McDowell et al., 2005), one measurement of the McClure apatite standard (523.51 ± 2.09 Ma, Schoene and Bowring, 2006), followed by six measurements of apatite crystals from the samples. This sequence was repeated four to six times, and the end of the session consisted on three analyses: one of the Durango and two of the Madagascar standards. Regarding the quality control and validation of the standards, the McClure apatite standard was dated at 528 ± 5 Ma (n = 31, MSWD = 0.5) in the first session and at 509 ± 12 Ma (n = 26, MSWD = 3) in the second session. The Durango standard was dated at 32 ± 0.5 Ma (n = 33, MSWD = 1.1) in the first session and at 30.6 ± 0.6 Ma (n = 16, MSWD = 1.1) in the second session. Further information on the dating protocol can be found in Pochon et al. (2016b) and in Supplementary material 3.

4 Results

4.1 Relationships between dolerites and Variscan schistosity cleavage

Regarding the structural data, the strike and dip of the contacts between the mafic intrusions and the host rocks were measured in the field (Fig. 4). Most of the mafic bodies in this study generally lie roughly parallel to the S0-1 cleavage developed in their host metasedimentary sequences (Fig. 4), suggesting that most of the intrusions are sills. The walls of the sills are sometimes affected by the 359–350 Ma (e.g. Dallmeyer et al., 1997) Variscan cleavage, whereas the cores are generally preserved. This observation demonstrates that dolerites were emplaced at least at the beginning of the Variscan orogeny.

thumbnail Fig. 4

Structural data of the mafic sills and metasedimentary sequence in the studied areas. Most of the mafic bodies in this study generally lie roughly parallel to the S0-1 cleavage developed in their host metasedimentary sequences.

4.2 Dolerite petrography

Samples from the CS present a doleritic texture and are mainly composed of Ca-plagioclase and clinopyroxene as main primary phases and chlorite as main secondary phase (Fig. 5A–F). Samples from GS and AS also present a doleritic texture and are mainly composed of primary plagioclase and secondary calcite and chlorite (Fig. 6A–F), with minor clinopyroxene (present in greater proportion in the AS rocks, compared to GS rocks, Fig. 6A and B). Other primary minerals include pyrrhotite, pentlandite, apatite and ilmenite. Ilmenite is less abundant than apatite and usually slightly altered to leucoxene. Accessory phases can also be found, either in the groundmass or locally filling some porosity, these include: leucoxene, chalcopyrite, zircon and baddeleyite (too small and too scarce to be dated). Secondary phases are ubiquitous and consist of actinolite, epidote, chlorite, calcite, and, in lesser amount, chalcopyrite and pyrite (Figs. 5 and 6). The dolerites from GS are generally more altered than those from the CS and AS. The least altered samples with low LOI (see Supplementary material 1 and 2) show primary features and abundant primary minerals compared to the samples with moderate to high alteration degree (Figs. 5 and 6). Samples with high LOI (see Supplementary material 1 and 2) show abundant secondary phases such as calcite and chlorite (Figs. 5 and 6). In most of the mafic rocks, apatite is relatively abundant and is present as acicular crystals within several primary mineral phases (e.g. plagioclase, clinopyroxene, ilmenite) as well as in the groundmass (Fig. 7).

thumbnail Fig. 5

Thin section micro-photographs showing some samples from the CS: A-B) Least altered sample with low LOI (SA2026) from CS, showing primary magmatic features and minerals such as plagioclase (Pl) and clinopyroxene (Cpx) and secondary phases (Chl-chlorite). C-D) Most altered sample (SA2042) from CS, most of the primary minerals are altered to secondary phases. Plagioclases and amphiboles are altered to epidote (Amph-Ep); E-F) Non-altered sample (SA2014B) with moderate LOI (5.8%) showing clinopyroxene (Cpx) as one of the primary phases and chlorite (Chl) as main secondary phase.
thumbnail Fig. 6

Thin section micro-photographs showing samples from AS and GS: A-B) Least altered sample with low LOI from AS (AL2088), showing primary features and phases such as plagioclase (Pl) and clinopyroxene (Cpx) with minor chlorite as main secondary phase; C-D) Altered sample from the GS (AL2075) with high LOI (10.5%), showing primary phases such as plagioclase and a big amount of secondary phases including chlorite (Chl) and calcite (Cal), which helps to explain the high LOI; E-F) Altered sample from AS (AL2077B) with high LOI (15.1%), showing primary phases such as plagioclase and secondary phases including chlorite (Chl) and calcite (Cal).
thumbnail Fig. 7

Selected thin section micro-photographs illustrating the studied apatite crystals. A) Optical image showing apatite associated to primary ilmenite; B) BSE image showing apatite associated to primary ilmenite; C-D) cathodoluminiscence micro-photographs of the studied apatite crystals, E-F) Thin section micro-photographs showing apatite related to primary clinopyroxene. The apatite appears as acicular crystals (Ap) associated to the primary plagioclases, clinopyroxene (Cpx) and ilmenite (Ilm).

4.3 Dolerites geochemistry

In order to quantify the alteration degree of the dolerites, we employed the Al-CCPI alteration box plot (Large, 2001), the AI-AAAI diagram (Williams and Davinson, 2004) and the isocon diagram (Grant, 1986, 2005) for metasomatic alteration (Figs. 8A, 8B, 9A and 9B). Most of the studied samples fall into the basalt field of the Al-CCPI alteration box plot (Fig. 8A), and into the unaltered box of the AI-AAAI diagram (Fig. 8B). The sample SA2022B falls within the altered rock field in both diagrams and will therefore not be considered for the remainder of this study (Fig. 8A, B).

The isocon diagram can be used to undertake a more detailed study of the alteration degree between lightly and heavily altered rocks, as well as the mobilization of major and trace elements (Grant, 1986, 2005). This diagram allows for the plotting of an altered composition against a non-altered (original) composition, thus facilitating the determination of its degree of alteration (Grant, 1986, 2005). Considering petrographical and whole rock geochemical analyses, we selected the least altered sample from CS (SA2026) and plotted the rest of the CS samples against it (Fig. 9A). In the case of the AS and GS, we used the least altered sample of the AS (AL2088) for all samples.

The isocon diagram for CS (Fig. 9A) shows that CaO, Fe2O3, Pb, Sr and LOI contents of altered rocks are higher, whereas SiO2, Na2O, U, Th and K contents are generally lower than those of the least altered sample. The two samples with the highest LOI also show the highest Pb contents (Fig. 9A), but the other samples do not show any significant correlation. The variations of MgO, SiO2, Al2O3 and P2O5 are relatively low, as well as those of REE, U, Th, Nb, Ta, Zr, Y, and generally Ti. On the contrary, Na2O, CaO, Fe2O3, K, Pb, Sr, Ba and Rb exhibit significant variations. AS samples show very similar features to CS samples, with the exception of LOI values that are substantially higher (Fig. 9B). Nonetheless, lower variations are observed for REE, U, Th, Nb, and Zr, even for AS samples with the highest LOI values (AL2077A and AL2077B, see Supplementary material 2). These results, together with the Al-CCPI alteration box plot (Large, 2001) and the AI-AAAI diagram (Williams and Davinson, 2004), allow us to determine that REE, U, Th, Nb, Ta, Zr, Y and Ti were probably unaffected by alteration processes or metamorphism and therefore can be used for petrogenetical interpretations of CS and AS samples, whereas K, Rb, Ba, Pb and Sr cannot.

The samples from the GS present high LOI, REE, Ba, P, and Mg contents as well as low Rb, Al, Na contents compared to AL2088. This may be due to an initially slightly enriched magmatic rock compared to AS and/or to later alteration. Impoverishments in Rb, K an Na can be also observed for these samples (Fig. 9B). The extremely high variability shown by GS samples for most elements (Fig. 9B) suggests that alteration or metamorphism may have affected the primary geochemical signature of these samples.

Most of the samples in this study plot into the basalt field, only two of the samples from the AS correspond to a basaltic andesite and basaltic trachyandesite (see TAS diagram, Le Bas et al., 1986, Supplementary material 4). However, we will make little use of this diagram since the proportions of alkalis may have been significantly affected by hydrothermal processes. According to the Nb/Y vs. Zr/Ti diagram (Pearce, 1996, 2008), the mafic sills from the AS and GS plot in the alkali basalt field, whereas those from the CS plot in the field of basalt (Fig. 10A). Using Zr/Y vs. Th/Yb classification diagram (Ross and Bédard, 2009), most of our CS samples plot between the transitional and the calc-alkaline field (Fig. 10B).

The intrusions display a slight enrichment in LREE with respect to HREE in chondrite normalized diagrams (Sun and McDonough, 1989; Figs. 11A and B). Samples from the CS show a regular REE pattern without Eu anomaly (Fig. 11A), whereas those from the AS and GS display a slight Eu negative anomaly and Lu enrichments (Fig. 11B). These samples also show higher La/Yb ratios (AS = 11.88; GS = 11.07) than those from the CS (3.27). Furthermore, the GS and AS rocks are enriched in LREE with respect to those from the CS. Primitive mantle normalized diagrams (values from Sun and McDonough, 1989) display high positive Pb anomalies whereas high negative Rb anomalies are observed in all dolerite samples. Slight positive Ti anomalies are observed for the dolerites from CS and AS and not for GS. Conversely, a slight positive Sr anomaly is observed only for the dolerites from AS. High negative K anomalies are observed in the dolerites from CS and GS whereas this anomaly is almost absent in the dolerites from AS (Fig. 11C and D).

thumbnail Fig. 8

Hydrothermal alteration diagrams for the samples in this study. A) Alteration box plot diagram of Large (2001). AI: Alteration Index, CCPI: Chlorite-Carbonate-Pyrite-Index; B) Alteration diagram of Williams and Davinson (2004), AAAI: Advanced-Argilic-Alteration-Index.
thumbnail Fig. 9

Isocon diagrams for metasomatic alteration of Grant (2005), showing samples from CS (A) and AS-GS (B). These diagrams allow to determine the alteration degree of all samples by plotting an altered composition against an original composition. A) Samples from CS plotted against the least altered sample from CS (SA2026); B) Samples from AS-GS plotted against the least altered sample from AS (AL2088), both synclines present the same stratigraphic succession and are found in the same region reason why GS samples where plotted against an AS sample.
thumbnail Fig. 10

Discrimination diagrams for the mafic samples of the CS, AS, and GS. A) Nb/Y vs. Zr/Ti diagram with field boundaries taken from Pearce (1996); B) Zr/Y vs. Th/Yb diagram with field boundaries taken from Ross and Bedard (2009).
thumbnail Fig. 11

A-B) Chondrite normalized spider diagram for samples from CS (A) and AS-GS (B); C-D) primitive mantle normalized spider diagrams for samples from CS (C) and the AS-GS (D). Primitive mantle normalizing values are from Sun and McDonough (1989).

4.4 Dolerites geochronology

4.4.1 Apatite features

Whatever the dolerite samples, apatite crystals consistently exhibit an acicular shape, with lengths ranging from 30 μm to 400 μm and widths between 20 μm and 60 μm. In cathodoluminescence images, the apatite crystals do not show any oscillatory zoning (Supplementary material 5). The lack of zoning along fractures suggests the lack of hydrothermal dissolution/recrystallization processes and supports its primary character (i.e. magmatic). The apatite grains exhibit homogeneous colours ranging from yellow, green, violet and blue.

Chemical compositions of the apatite grains from the seven dated dolerite samples are shown in Supplementary material 6. A total of 25–50 apatite crystals were analyzed for each rock sample, depending on the quantity and size of the available apatite. To examine the chemistry of apatite, EPMA data were plotted in the OH-Cl-F diagram (Supplementary data 6). All data lies near the F apex, highlighting fluorapatite compositions in line with F contents exceeding 0.62 apfu (Supplementary material 7). Apatite grains from the Almaden dolerite (AL2076) are the closest to the fluorapatite pole. These fluorapatite composition match quite well that of apatite crystals extracted from the Armorican dolerites (green field in Supplementary material 7). Finally, the Grant diagram (Fig. 9) reveals only limited variations in phosphorus content, which suggests that apatite remained largely unaffected by post-magmatic processes.

4.4.2 U-Pb dating on apatite

The U-Pb data of apatite from seven samples, three from CS (SA2014A, SA2022C and SA2026), three from GS (AL2072, AL2073 and AL2075) and one from AS (AL2076) (Figs. 12 and 13) are presented in Tera-Wasserburg diagrams made using IsoplotR (Vermeesch, 2018). All errors are provided at 2s (%) level, as shown with ellipses in Figures 11, 12 and in Supplementary material 8. The rest of geochronological data for the dated samples can be found in Supplementary material 8.

Sills from CS yielded 3 dates (Fig. 12). Sample SA2014A (Fig. 12A) yielded a lower intercept date of 359 ± 21 Ma (MSWD = 1.5; n = 34) with an initial 207Pb/206Pb value of 0.834 ± 0.077. Sample SA2022C (Fig. 12B) provided a lower intercept date of 353 ± 19 Ma (MSWD = 0.95; n = 35) with an initial 207Pb/206Pb value of 0.857 ± 0.034. For sample SA2026 (Fig. 12C), the lower intercept is 332 ± 17 Ma (MSWD = 0.97; n = 58) with an initial 207Pb/206Pb value of 0.819 ± 0.039.

The rocks of the GS yielded 3 dates. The sample AL2075 (Fig. 13, GS) provided a lower intercept date of 422 ± 11 Ma (MSWD = 1.1; n = 33) with an initial 207Pb/206Pb value of 0.828 ± 0.010. For the sample AL2072 (Fig. 13B, GS), the lower intercept is 387 ± 18 Ma (MSWD = 0.92; n = 21) with an initial 207Pb/206Pb value of 0.812 ± 0.033. Sample AL2073 (Fig. 12C, GS) yielded a lower intercept date of 361 ± 22 Ma (MSWD = 2, n = 34) with an initial 207Pb/206Pb value of 0.796 ± 0.054.

Finally, sample AL2076 (Fig. 13D, AS) provided a lower intercept date of 358 ± 12 Ma (MSWD = 0.92; n = 34) with an initial 207Pb/206Pb value of 0.818 ± 0.020.

thumbnail Fig. 12

Tera-Wasserbourg diagram showing the obtained dates for the mafic rocks of the CS.
thumbnail Fig. 13

Tera-Wasserbourg diagram showing the obtained dates for the mafic rocks coming from GS (A-C) and AS (D).

5 Discussion

5.1 Interpretation of the apatite dating and significance of the ages

As shown in Figure 7, the apatite crystals that were dated in this study present magmatic textures and are associated to primary magmatic phases (plagioclase, clinopyroxene). Furthermore, the mineral chemistry of the dated crystals in the CS-AS-GS (Supplementary material 6 and 7) is typical of igneous apatite in mafic rocks (Kieffer et al., 2024) and matches with the magmatic apatites dated at ca. 360 Ma in the Armorican Massif (e.g. Pochon et al. 2016b). The low-grade metamorphism affecting the whole CS area explains the partial replacement of the magmatic minerals in the dolerites by chlorite, green amphibole, albite, calcite and actinolite, which are minerals characteristic of the greenschist facies (López-Moro et al., 2007, 2020). In the case of the Almadén district, Higueras et al. (1995) reported metamorphic conditions that reached a maximum temperature of 240 °C and pressure around 200 MPa during the Variscan orogeny. Similar conditions, corresponding to the transition between anchizonal and greenschist metamorphic facies, are reported for the studied areas in the CIZ by Martínez Catalán et al. (2014). Such a low-grade metamorphism is unlikely to have reset the apatite isotopic system (Cherniak et al., 1991; Pochon et al., 2016b). The closure temperature of the isotope system of this mineral commonly ranges between 375 and 550 °C (Dodson, 1973; Webster and Piccoli, 2015), although recent studies determined higher closing temperatures ranging between 770 and 870°C (e.g. Pochon et al. 2016b). These temperatures were based on the fast cooling rate and low thickness of several mafic sills in the Armorican Massif (Pochon et al. 2016b), conditions that are similar to those observed for the sills studied in this work. All these observations support the fact that the measured dates can be interpreted as crystallization ages.

5.1.1 -Significance of ages in the Guadalmez and Almadén synclines

The numerous alkaline dolerite sills in the GS intruded mainly black schists that have been ascribed to Silurian, ca. 420–440 Ma, based on their stratigraphic position and graptolite fossils (Almela et al., 1962; García San Segundo et al., 1985; Lorenzo Álvarez et al., 1995; Villas et al., 1999). The three dolerite sills dated in this study provided ages of 422 ± 11 Ma (AL2075), 387 ± 18 Ma (AL2072), and 361 ± 22 Ma (AL2073). The oldest age could be coeval with the host sedimentary rocks (within uncertainty), which means that the dolerites could be formed during or shortly after active sedimentation. In the central part of the GS, sedimentary rocks with Carboniferous ages (ca. 360-300 Ma; see Fig. 3) were described (Almela et al., 1962; García San Segundo et al., 1985; Lorenzo Álvarez et al., 1995), which means that the sedimentation in the area continued after the intrusion of the dolerites. The lack of overlap between the age of the oldest sill and the younger ones suggests at least two different magmatic events in the GS (Fig. 14).

The alkaline dolerite sills in the AS intrude Upper Devonian (ca. 385–370 Ma) volcano-sedimentary rocks (Almela et al., 1962; García San Segundo et al., 1985; Lorenzo Álvarez et al., 1995; Villas et al., 1999; Fig. 3). The dated dolerite sample AL2076 yielded an age of 358 ± 12 Ma, which is coeval with the host rock within error, and which indicates that it was intruded during or shortly after the deposition of the volcano-sedimentary unit. This age matches the ages found for the younger dolerites dated in the GS (ca. 15km SW from the AS), which could mean that they were part of the same magmatic event (Fig. 14).

thumbnail Fig. 14

Synthesis of the ages of the mafic samples studied in this work and comparison with previous data on mafic magmatism in the Armorican Massif. References in the figure: 1(Pochon et al., 2016b); 2(Barboni et al., 2013); 3(Paquette et al., 2017).

5.1.2 Significance of the ages in the Codosera syncline

In the CS, the dolerite sills intrude the Gévora Formation, which was deposited during the Carboniferous Period (Mississippian, ca. 359–323 Ma), an age estimation that is based on the presence of miospores in the black schist and a conodont assemblage in the San Antonio limestone (Rodríguez González et al., 2007; Fig. 3). The studied dolerites in the CS yielded ages of 359 ± 21 Ma (SA2014A), 353 ± 19 Ma (SA2022C) and 332 ± 17 Ma (SA2026), suggesting that the dolerite sills were emplaced during or shortly after the sedimentation. Unfortunately, the uncertainties of the measured ages from the mafic sills of the CS are too large to conclude whether the sills were formed during a long magmatic event or as a sequence of smaller events, as in the GS (Fig. 14). The ages of the CS samples are within uncertainty similar to the dolerites of the AS and the younger dolerites of the GS (Fig. 14), which are separated ca. 200km SE from the CS.

5.1.3 Magmatic events in the GS, AS and CS over ≈120 Ma

If we synthesise the geochronological data obtained in this work, the emplacement ages of the dolerite sills of the GS, AS and CS spread out over 120 Ma, considering the associated uncertainty (Fig. 14). Two main types of magmatism are identified, one of alkaline affinity in the GS and AS, and the second of transitional affinity in the CS (Fig. 9). Considering the spatial distribution of the dolerites, together with the geochemistry and geochronological data, we propose the existence of at least three mafic magmatic events, two in the GS-AS and one in the CS: i) the first one of alkaline affinity during the upper Silurian (i.e. 420 Ma) in the GS (Fig. 14); ii) a second event(s) of alkaline affinity in the GS and AS, taking place either as a single long event at about 370±34 Ma or as short multiple events between 387 ± 18 Ma and 360 ± 22 Ma (Fig. 14); iii) a third event(s) in the CS, showing a transitional affinity, which took place as a single major event at about 348±33Ma or as a series of shorter events from 359 ± 21 Ma down to 332 ± 17 Ma (Fig. 14). More precise dating of the dolerites could reveal if these magmatic events were characterized by short-lived syncline-scale magma pulses or sequential intrusions of smaller sills at different parts of the synclines.

5.2 Major geological events, evolution and geochemistry of the mafic magmatism in the CIZ

The mafic magmatism in the CIZ has experienced an important geochemical evolution from the first stages of the Palaeozoic era until the end of the Variscan orogeny, the signatures of all of these rocks being strongly correlated to their geodynamic setting.

5.2.1 Silurian to Lower Devonian

The opening of the Rheic ocean stopped during Silurian to Lower Devonian (ca. 430–415 Ma; Arenas et al., 2007; Martínez Catalan et al., 2007; Nance et al., 2012) reaching a maximum width of ca. 4000 km (Nance et al., 2012 and references therein). The age of the oldest dolerite in this study (AL2075, GS, alkaline affinity; 422 ± 11 Ma) is coeval with the last stage of this extensive period and, within uncertainty, with the depositions of its host sedimentary rocks. This sample shows an enriched metasomatized mantle source possibly related to the Cadomian slab as in the case of the Carrascal Massif (e.g. Solá et al., 2010). However, there is no clear sign of deep crustal recycling in this mafic sample from the GS.

5.2.2 Lower Devonian to Carboniferous

The subduction of the Rheic ocean started at ca. 415 Ma and lasted until its closure at the Variscan collision during Upper Devonian to early Carboniferous (Arenas et al., 2007; Martínez Catalan et al., 2007; Nance et al., 2012). In the CIZ, the first collisional stage of the Variscan orogeny has been dated at 359–350 Ma (Dallmeyer et al., 1997). In this geodynamic setting, the Badajoz-Córdoba shear zone accommodated the deformation by left lateral transtensional kinematics, leading to an extensional event documented at the southern part of the CIZ (Martínez-Catalán et al., 2016; Azor et al., 2019; Gutiérrez-Marco et al., 2019).

This geodynamical setting facilitated the intrusion of several mafic suites across the CIZ, which correspond to: i) the alkaline basalts from El Castillo (394.7 ± 1.4 Ma, Gutiérrez-Alonso et al., 2008); ii) the alkaline mafic rocks from the GS and AS in this study (emplaced between 387 ± 18 Ma to 360 ± 22 Ma) as well as those from the AS reported in previous studies (e.g. Hall et al., 1997; Higueras, 1995, 2000; Higueras et al., 2013); iii) the transitional dolerite sills from the CS reported in this study that were emplaced at 359 ± 21 Ma, 353 ± 19 Ma and 332 ± 17 Ma, and iv) the mafic and ultramafic intrusions documented in the IBERSEIS Reflective Body (IRB; Simancas et al., 2003, 2006), which can be observed in the middle crust of the southern CIZ, the Ossa Morena Zone and the South Portuguese Zone (Simancas et al., 2003, 2006). All of them have been related to the transtensional kinematics of the Badajoz-Cordoba shear zone at the first collisional stage of the Variscan orogeny (Simancas et al., 2003, 2006; Carbonell et al., 2004; Palomeras et al., 2011; Cambeses et al., 2015; Azor et al., 2019).

5.2.3 Possible sources and geochemical signatures

5.2.3.1 The Almadén and Guadalmez synclines

Regarding the geochemistry of the GS and AS rocks in this study, the correlation between Nb/Yb and Th/Yb (Pearce, 2008) suggests a geochemical signature related to an enriched mantle source without crustal contamination (Fig. 15). Furthermore, all of the mafic rocks from the GS and the AS show REE patterns with low HREE/LREE/ (Fig. 11B and D), a feature that has been usually related with the presence of residual garnet in the source (Fig. 16), because this mineral is capable of concentrating HREE in its structure (Harrison, 1981; Kamenetsky and Eggins, 2012). Mafic rocks presenting high Gd/Yb (>1.5) and high Ti/Y (>350) ratios, such as the samples from GS (Gd/Yb>3; Ti/Y>559) and AS (Gd/Yb>2.9; Ti/Y>554), may also indicate the presence of residual garnet in the source (e.g. Song et al., 2009; Kamenetsky and Eggins, 2012; Sarrionandia et al., 2023; Fig. 15A).

The negative anomalies in K, Rb and Ba observed for the GS and AS samples (Fig. 11B and D) were also reported in the previously studied mafic rocks in the AS (Higueras, 1995; Higueras et al., 2013). Considering the alteration studies performed in section 4.3, these anomalies may be attributed to alteration or metamorphic processes.

We suggest that the mantle source for the mafic rocks in the GS and AS is characterized mostly by the dry garnet-Cpx assemblage, which could form in the Sub Continental Lithospheric Mantle (SCLM) as a result of dry metasomatism (Griffin et al., 1999).

The seismic studies and 3D velocity models of Palomeras et al. (2017) show low velocity zones that we could interpret as metasomatized SCLM possibly influenced by a Cadomian/Rheic slab, which has been suggested as the mantle source of several mafic suites in the OMZ and CIZ (Sarrionandia et al., 2023).

thumbnail Fig. 15

Th/Yb vs. Ta/Yb diagram from Pearce (1983, 2008) showing the mafic rocks studied in this work and those reported for the CIZ and the Armorican Massif. The rocks from CS show a less enriched and less metasomatized mantle source compared to the rocks from AS and GS. The CS is consistent with the data reported for the Armorican Massif (Pochon et al., 2016b). References in the figure: 1(Orejana et al., 2009); 2(Scarrow et al., 2009); 3(López-Moro et al., 2007); 4(Pochon et al., 2016b).
thumbnail Fig. 16

A) La/Sm vs. Gd/Yb diagram (Kamenetsky and Eggins, 2012) showing samples coming from La Codosera, Almadén and Guadalmez synclines. The data is consistent with the presence of garnet in the source, which also matches with the enrichment in LREE.
5.2.3.2 Codosera syncline

In the case of the CS, the magmatic suites reported in this study show a subalkaline-transitional affinity (samples SA2014A-359 ± 21 Ma, SA2022C-353 ± 19 Ma and SA2026-332 ± 17 Ma). Previous studies suggested that the mafic magmas in the CS presented a subalkaline affinity and were derived from a spinel-bearing/garnet-free metasomatized sub lithospheric mantle source (López-Moro et al., 2007). Nevertheless, the trace element systematics of our samples are not compatible with this last hypothesis. The Gd/Yb (1.98–3.10) and Ti/Y (415 to 700) ratios of the dolerites in the CS are geochemically similar to those of the dolerites in the AS and GS, as well as of other mafic magmas derived from garnet-bearing mantle sources, in the Iberian Massif and across the globe (e.g. Song et al., 2009; Kamenetsky and Eggins, 2012; Sarrionandia et al., 2023). The presence of residual garnet in the mantle source is also supported by the Gd/Y vs. La/Sm diagram from Kamenetsky and Eggins (2012), where all data plot in the garnet-source field but show lower amount of garnet in the source compared to AS and GS (Fig. 16). The lesser LREE contents of the CS (compared to those from the GS and AS), and the lower La/Yb ratios (3.27 versus 11.88 for AS samples and 11.07 for GS ones) suggest higher degrees of partial melting to produce the CS rocks (Fig. 11A and B). Our data is consistent with previous studies suggesting 10-35% of partial melting for the CS rocks (López-Moro et al., 2007; Villaseca et al., 2022), and ≈5% partial melting for the AS mafic suites (Higueras, 1995; Higueras et al., 2013; Villaseca et al., 2022).

The mafic rocks from the CS reported by López-Moro et al. (2007) and most of the samples in this work, lie in the MORB-OIB array (Fig. 15), and indicate an enriched mantle source with lower Nb/Yb and Th/Yb ratios, compared to those from the AS and GS (Fig. 14). This clear difference in the mantle source suggests that the magmatism in the CS is a separate event compared to the AS and GS, although their ages are overlapping within uncertainty. Crustal interaction seems to be of minor importance in all studied areas compared to the rest of the mafic suites in the SCS (Fig. 15) emplaced during the last stages of the Variscan orogeny (Scarrow et al., 2009; Orejana et al., 2017).

Similarly to GS and AS samples, CS rocks show pronounced negative anomalies in K and Rb, and to a lesser extent in Ba, which were already observed in the CS rocks and attributed to hydrothermal and alteration effects (López-Moro et al., 2007).

5.3 The Variscan geodynamic setting: The emplacement of mafic magmas in the Armorican Massif and the CIZ

Some of the ages of the dolerites in this study are coeval with the emplacement of mafic intrusions in the Armorican Massif, roughly 1000 km to the NE from the CS, GS and AS (Figs. 1 and 14; Pochon et al., 2016b, 2019). Based on this temporal correlation, we propose the existence of a widespread mafic magmatism that took place from the Upper Devonian to the Lower Mississippian in the CIZ and in the Armorican Massif. The rocks from the Armorican Massif present an enriched metasomatized mantle source with Nb/Yb-Th/Yb ratios similar to CS, and lower compared to the rocks from the AS and GS (Fig. 15). The mineral chemistry of the dated dolerites in the Armorican Massif (e.g. magmatic apatites; Pochon et al. 2016b) matches with the one documented for the CIZ mafic rocks in this study.

Regarding the geodynamic setting at large scale, the emplacement of the dolerites in the Armorican Massif was associated with a slab roll back, followed by a large scale astenospheric upwelling, which led to the partial melting of the Sub Continental Lithospheric Mantle, causing a crustal thinning and an extensional event in the external zone of the Variscan Belt (Central and North Armorican Domain) during a convergent geodynamic setting (Pochon, 2018). In this particular context, the Variscan front was situated in the future South Armorican Domain, a considerable distance (≈100km) from the domains in which dolerites were emplaced (Pochon, 2018).

In the case of the CIZ of the Iberian Massif, the geodynamic setting associated with the emplacement of the dolerites in this study is related to the opening (GS; 422 ± 11 Ma; Fig. 17A), subduction (GS, this study, 387 ± 18 Ma; Basalts El Castillo-EC-394.7 ± 1.4 Ma, Gutiérrez-Alonso et al., 2008; Fig. 17B) and closing of the Rheic ocean/Variscan collision (AS-GS-CS, 361 ± 22 Ma–332 ± 17 Ma; Fig. 17C). During this large scale collisional event, a slab break-off is proposed for explaining the emplacement of large volumes of mafic magmas in the Ossa-Morena-Zone (e.g. Beja Igneous Complex; Pin et al., 2008) and the CIZ (e.g. IBERSEIS Reflective Body; Simancas et al., 2003, 2006). This slab break-off hypothesis, may also have contributed to the emplacement of some of the mafic intrusions reported in this study in the CS and AS-GS, which are located ≈100km from the Variscan front (Fig. 17C).

Some of the ages reported in this study are also concordant with a Sb peak mineralization at about 360 Ma in the Armorican Massif (e.g. Pochon et al., 2019), which is also coeval with the emplacement of mafic intrusions in this domain (e.g. Pochon et al., 2016b). It has been suggested that the concordance between the ages of the mafic events and those of Sb mineralization imply a possible genetical link between mafic magmatism and ore forming processes in the Armorican Massif (Pochon, 2018; Pochon et al., 2018, 2019). In the CS, the hydrothermal event associated with mafic volcanism was proposed as a possible source for the mineralizing fluids leading to the formation of the Sb deposits in the area (Arribas and Gumiel, 1984). However, the lack of geochronological data for the Sb deposits makes the genetical link between Sb mineralization and the mafic magmatism difficult to prove. Further studies are needed in order to understand if the Sb deposits in the CS are coeval with the mafic rocks dated in this study in the same way than in the Armorican Massif.

As mentioned above, several Hg deposits have been identified in the AS. The formation of some of these deposits has been dated between 427 and 375 Ma, suggesting that they were generated in response to several mineralizing events over a period of 50 M.y (Hall et al., 1997; Palero and Lorenzo, 2009; Higueras et al., 2013; Palero-Fernández et al., 2015). Based on field studies, it was suggested that mafic volcanism was coeval with the Hg mineralization, possibly acting as the trigger for the hydrothermal activity leading to the formation of these deposits (Higueras, 1995; Hernández et al., 1999; Higueras et al., 2013). Our geochronological data confirms that the mafic magmatism was indeed active in the AS and GS during this period.

thumbnail Fig. 17

Large scale geodynamic setting model proposed for the emplacement of the mafic rocks in the GS-AS-CS. This figure was created and modified after the tectonic evolution models made by Simancas et al. (2003, 2006); Gutiérrez-Alonso et al. (2008); Pin et al. (2008); Martínez Catalán et al. (2009). A) Opening of the Rheic ocean (GS; 422 ± 11 Ma); B) subduction of the Rheic ocean (GS, this study, 387 ± 18 Ma; Basalts El Castillo-EC-394.7 ± 1.4 Ma, Gutiérrez-Alonso et al., 2008; Fig. 17B) and; C) closing of the Rheic ocean/Variscan collision (AS-GS-CS, 361± 22 Ma-332 ± 17 Ma; Beja Igneous Complex-BIC; Pin et al., 2008).

6 Conclusions

The LA-ICP-MS dating of apatite, together with whole-rock geochemistry analyses helps in the understanding of the geochronological and geochemical characteristics of several intrusive mafic rocks across three different domains of the CIZ (GS, AS and CS). The alkaline dolerites from the GS and AS were emplaced during at least two distinct events, the first one occurring approximately 420 Ma ago and the second between 387 ± 18 Ma and 360 ± 22 Ma. The dolerites from the CS were emplaced between 359 ± 21 Ma and 332 ± 17 Ma and are related to a magmatism of transitional affinity. The Silurian to Lower Devonian age in the GS is consistent with the final stage of the opening of the Rheic ocean. The Lower Devonian to Carboniferous ages in the GS, AS and CS span from the subduction/closure of the Rheic ocean to the onset of the Variscan orogeny. This magmatism is related to local extensional settings, which have been documented at the southern part of the CIZ during the initial setting of the Variscan orogeny. The AS and GS rocks are also coeval with Hg mineralization events documented in the AS during Silurian and Devonian periods (Higueras, 1995; Higueras and Oyarzun, 2005; Higueras et al., 2013; Palero-Fernández et al., 2015).

The Upper Devonian to early Carboniferous mafic suites in the GS, AS, and CS coincide with the emplacement ages of several mafic rocks in the Armorican Massif (Pochon et al., 2016b), allowing us to propose that a widespread mafic magmatism took place in both domains (CIZ and Armorican Massif) at ca. 360-350 Ma.

According to this study, the middle Silurian to Upper Devonian alkaline magmatism in the AS and GS may be related to a garnet bearing enriched mantle source with no evidence of crustal recycling or contamination and a low degree of partial melting. The Upper Devonian to Carboniferous mafic magmatism in the CS seems to be associated with a less enriched, garnet-bearing source, similar to the mafic suites of the Armorican Massif (Pochon et al., 2016b). Furthermore, the lower LREE/HREE ratios of the CS compared to GS and AS might imply the highest degree of partial melting among all of the studied samples in this study. Additional geochemical analyses (e.g. isotope geochemistry) are crucial in order to better constrain the sources and the evolution of the Upper Devonian-lower Carboniferous mafic magmatism in the CIZ.

Acknowledgments

This article is respectfully dedicated to the memory of Dr. Eric Gloaguen (BRGM), who directed the AUREOLE project from its inception until his premature demise. This work was funded by the ANR (ANR-19-MIN2-0002-01), the AEI (MICIU/AEI/REF.: PCI2019-103779) and author’s institutions in the framework of the ERA-MIN2 AUREOLE project (i.e., tArgeting eU cRitical mEtals (Sb, W) and predictibility of Sb-As-Hg envirOnmentaL issuEs; https://aureole.brgm.fr), as well as by LabEx VOLTAIRE (ANR-10-LABX-100-01), EquipEx PLANEX (ANR-11-EQPX- 0036) and Project SBPLY/17/180501/000273, Consejería de Educación, Regional Government of Castilla-La Mancha, Spain. We thank the company Villar del Rey Natural Stones in the Extremadura Region for their support by letting us work and sample in their quarries. The authors appreciate the valuable discussions with Dr. David Orejana García (Universidad Complutense de Madrid, Spain) Dr. Jérémie Melleton (BRGM, France) and Dr. Bryan Cochelin (ISTO, France).

Supplementary Material

Supplementary material 1. Whole-rock geochemistry data for studied mafic samples from la Codosera syncline-Central Iberian Zone.

Supplementary material 2. Whole-rock geochemistry data for studied mafic samples from Guadalmez and Almadén synclines-Central Iberian Zone.

Supplementary material 3. Operating conditions for the LA-ICP-MS equipment

Supplementary material 4. Total Alkali Silica diagram (TAS) from Le Bas et al. (1986) showing the samples studied in this work and in previous studies done by Higueras (1995); Higueras et al. (2013); López-Moro et al. (2007).

Supplementary material 5. Cathodoluminescence micro-photographs of the studied apatite crystals for samples A) SA2022C; B) SA2026; C) AL2075; D) AL2076. In cathodoluminescence images, the apatite crystals do not show any oscillatory zoning. The lack of any zoning along fractures suggests the lack of hydrothermal dissolution/recrystallization processes and supports their main primary character. The apatites exhibit homogeneous colours ranging from yellow, green, violet and blue

Supplementary material 6. Average electron microprobe analyses and corresponding structural formulas of apatite for the mafic samples dated in this study.

Supplementary material 7. Average electron microprobe analyses of apatite for the dolerites samples dated in this study, compared to the apatites from the Armorican massif dolerites (field in green) by Pochon et al. (2016b). The halogens content plot has been modified from Boudreau (1995).

Supplementary material 8. Representative LA-ICP-MS U-Pb apatite data for studied samples.

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Cite this article as: Campos-Rodríguez HR, Gloaguen E, Tuduri J, Pochon A, Iacono-Marziano G, Poujol M, Higueras P, Virtanen VJ, Sizaret S, Lorenzo S, Esbri J-M, Mollé V. 2025. Geochemistry and geochronology of the Palaeozoic mafic magmatism in the Codosera, Almadén and Guadalmez synclines, Central Iberian Zone, Spain, BSGF - Earth Sciences Bulletin 196: 15. https://doi.org/10.1051/bsgf/2025008

All Figures

thumbnail Fig. 1

A) Geological sketch of the Variscan Belt showing the main terranes and domains, modified after Martínez Catalán et al. (2014); B) Simplified geological map of the Iberian Massif (Julivert et al., 1974; Simancas, 2019) showing mafic rocks in the studied areas and the mafic suites reported in the Central Iberian Zone (CIZ) from Ordovician times to the late stages of the Variscan orogeny. Abbreviations: OMZ: Ossa Morena Zone; GTMZ: Galicia-Tras-Os Montes-Zone; WALZ: West Asturian Leonese Zone (WALZ); CZ: Cantabrian Zone; MN: Montagne Noire; ECM: External Crystalline Massifs of the Alps; VM: Vosges Massif; BF: Black Forest; CP: Calzadilla Plutón (López-Moro et al., 2017); LA: Praia de Labruge (de Almeida Coelho, 2021); GD: Godomar (Medeiros et al., 1981); F: Farminhao (Cotrim et al., 2021); AP: Appinites (Montero et al., 2004b, a; Scarrow et al., 2009); TAC: Toledo Anatetic Complex (Bea et al., 2006); C: Caloco Sector (Villaseca et al., 2015; Orejana et al., 2017); R: Revenga Sector (Villaseca et al., 2015; Orejana et al., 2017); TM: Tenzuela Massif; MMC: Mérida Montoro Complex(Bandrés et al., 2004); SB: Sierra Bermeja (Errandonea-Martin et al., 2019); CM: Carrascal Massif (Solá et al., 2010); MG : Microgabbros (Orejana et al., 2020); CS: Codosera Syncline (López-Moro et al., 2007); AS: Almadén Syncline (Higueras, 1995; Higueras et al., 2013); GS: Guadalmez Syncline (Gumiel and Arribas, 1987; Lorenzo Álvarez et al., 1995).
In the text
thumbnail Fig. 2

A) Simplified geological map of the Iberian Massif showing the studied area of La Codosera syncline (CS). Modified after López-Moro et al. (2020) and references therein ; B) Geological map of the CS showing the mafic samples studied in this work. This map was created based on the Geological Maps (scale 1:50000) of Alburquerque (Santos García and Casas Ruiz, 1982), Puebla de Obando (López Díaz et al., 1989), Villar del Rey (Santos García et al., 1990) and Botoa (Insua Marquez et al., 1990).
In the text
thumbnail Fig. 3

A) Simplified geological map of the Iberian Massif showing the studied area of the Almadén (AS) and Guadalmez (GS) synclines. Modified after López-Moro et al. (2020) and references therein ; B) Geological map of the of the Almadén (AS) and Guadalmez (GS) synclines showing the mafic samples studied in this work. This map was created based on the Geological Maps (scale 1:50000) of Chillón (Lorenzo Álvarez et al., 1995), Almadén (García San Segundo et al., 1985), Hinojosa del Duque (Mira López et al., 1983a) and San Benito (Mira López et al., 1983b).
In the text
thumbnail Fig. 4

Structural data of the mafic sills and metasedimentary sequence in the studied areas. Most of the mafic bodies in this study generally lie roughly parallel to the S0-1 cleavage developed in their host metasedimentary sequences.
In the text
thumbnail Fig. 5

Thin section micro-photographs showing some samples from the CS: A-B) Least altered sample with low LOI (SA2026) from CS, showing primary magmatic features and minerals such as plagioclase (Pl) and clinopyroxene (Cpx) and secondary phases (Chl-chlorite). C-D) Most altered sample (SA2042) from CS, most of the primary minerals are altered to secondary phases. Plagioclases and amphiboles are altered to epidote (Amph-Ep); E-F) Non-altered sample (SA2014B) with moderate LOI (5.8%) showing clinopyroxene (Cpx) as one of the primary phases and chlorite (Chl) as main secondary phase.
In the text
thumbnail Fig. 6

Thin section micro-photographs showing samples from AS and GS: A-B) Least altered sample with low LOI from AS (AL2088), showing primary features and phases such as plagioclase (Pl) and clinopyroxene (Cpx) with minor chlorite as main secondary phase; C-D) Altered sample from the GS (AL2075) with high LOI (10.5%), showing primary phases such as plagioclase and a big amount of secondary phases including chlorite (Chl) and calcite (Cal), which helps to explain the high LOI; E-F) Altered sample from AS (AL2077B) with high LOI (15.1%), showing primary phases such as plagioclase and secondary phases including chlorite (Chl) and calcite (Cal).
In the text
thumbnail Fig. 7

Selected thin section micro-photographs illustrating the studied apatite crystals. A) Optical image showing apatite associated to primary ilmenite; B) BSE image showing apatite associated to primary ilmenite; C-D) cathodoluminiscence micro-photographs of the studied apatite crystals, E-F) Thin section micro-photographs showing apatite related to primary clinopyroxene. The apatite appears as acicular crystals (Ap) associated to the primary plagioclases, clinopyroxene (Cpx) and ilmenite (Ilm).
In the text
thumbnail Fig. 8

Hydrothermal alteration diagrams for the samples in this study. A) Alteration box plot diagram of Large (2001). AI: Alteration Index, CCPI: Chlorite-Carbonate-Pyrite-Index; B) Alteration diagram of Williams and Davinson (2004), AAAI: Advanced-Argilic-Alteration-Index.
In the text
thumbnail Fig. 9

Isocon diagrams for metasomatic alteration of Grant (2005), showing samples from CS (A) and AS-GS (B). These diagrams allow to determine the alteration degree of all samples by plotting an altered composition against an original composition. A) Samples from CS plotted against the least altered sample from CS (SA2026); B) Samples from AS-GS plotted against the least altered sample from AS (AL2088), both synclines present the same stratigraphic succession and are found in the same region reason why GS samples where plotted against an AS sample.
In the text
thumbnail Fig. 10

Discrimination diagrams for the mafic samples of the CS, AS, and GS. A) Nb/Y vs. Zr/Ti diagram with field boundaries taken from Pearce (1996); B) Zr/Y vs. Th/Yb diagram with field boundaries taken from Ross and Bedard (2009).
In the text
thumbnail Fig. 11

A-B) Chondrite normalized spider diagram for samples from CS (A) and AS-GS (B); C-D) primitive mantle normalized spider diagrams for samples from CS (C) and the AS-GS (D). Primitive mantle normalizing values are from Sun and McDonough (1989).
In the text
thumbnail Fig. 12

Tera-Wasserbourg diagram showing the obtained dates for the mafic rocks of the CS.
In the text
thumbnail Fig. 13

Tera-Wasserbourg diagram showing the obtained dates for the mafic rocks coming from GS (A-C) and AS (D).
In the text
thumbnail Fig. 14

Synthesis of the ages of the mafic samples studied in this work and comparison with previous data on mafic magmatism in the Armorican Massif. References in the figure: 1(Pochon et al., 2016b); 2(Barboni et al., 2013); 3(Paquette et al., 2017).
In the text
thumbnail Fig. 15

Th/Yb vs. Ta/Yb diagram from Pearce (1983, 2008) showing the mafic rocks studied in this work and those reported for the CIZ and the Armorican Massif. The rocks from CS show a less enriched and less metasomatized mantle source compared to the rocks from AS and GS. The CS is consistent with the data reported for the Armorican Massif (Pochon et al., 2016b). References in the figure: 1(Orejana et al., 2009); 2(Scarrow et al., 2009); 3(López-Moro et al., 2007); 4(Pochon et al., 2016b).
In the text
thumbnail Fig. 16

A) La/Sm vs. Gd/Yb diagram (Kamenetsky and Eggins, 2012) showing samples coming from La Codosera, Almadén and Guadalmez synclines. The data is consistent with the presence of garnet in the source, which also matches with the enrichment in LREE.
In the text
thumbnail Fig. 17

Large scale geodynamic setting model proposed for the emplacement of the mafic rocks in the GS-AS-CS. This figure was created and modified after the tectonic evolution models made by Simancas et al. (2003, 2006); Gutiérrez-Alonso et al. (2008); Pin et al. (2008); Martínez Catalán et al. (2009). A) Opening of the Rheic ocean (GS; 422 ± 11 Ma); B) subduction of the Rheic ocean (GS, this study, 387 ± 18 Ma; Basalts El Castillo-EC-394.7 ± 1.4 Ma, Gutiérrez-Alonso et al., 2008; Fig. 17B) and; C) closing of the Rheic ocean/Variscan collision (AS-GS-CS, 361± 22 Ma-332 ± 17 Ma; Beja Igneous Complex-BIC; Pin et al., 2008).
In the text

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