Volcanic activity recorded over time is often displayed in ancient marine environments as multi-episodic volcanogenic deposits interbedded within sediments. Dating these volcanic manifestations and ascertaining the genuine age of the volcanism are essential steps with a direct impact on understanding the geodynamic context of the sedimentary basin, as well as a significant starting point for many different studies (e.g., magnetostratigraphic ones, including the refinement of the Geomagnetic Polarity Time Scale for periods preceding the present-day oceanic magnetic record).
In this work, a case study of a shallow-marine and exceptionally ammonite- and brachiopod-rich Lower–Middle Jurassic carbonate succession of the Iberian Range (Spain) is presented. They were deposited along the Iberian shelf during the opening of the westernmost Tethys Ocean. These carbonate sediments include a series of volcanic levels mostly made of explosive pyroclastic to epiclastic deposits and rare lava flows, whose mineral composition shows alkaline affinity compatible with an extensional regime [e.g., Gautier 1968; Ortí and Sanfeliu 1971; Gómez 1979; Ortí and Vaquer 1980; Ortí 1987; Ancochea et al. 1988; Lago et al. 1996, 2004; Martínez et al. 1996a; Martínez et al. 1996c, d; Martínez et al. 1997a; Martínez et al. 1998; Valenzuela et al. 1996; Cortés 2018].
The marine strata containing the interbedded volcanic levels display a high fossil content (particularly ammonites and brachiopods) with possibilities of providing accurate biostratigraphic ages. This is possible because of the rapid evolution over time of ammonoids, allowing near-global precise standard zonations [e.g., Torrens 2002; Page 2003; Callomon 2003].
This scenario would be the reverse situation of the “interbedded volcanic rocks (PLAN A)” of Copeland , in which fossiliferous-rich sedimentary strata are located above and below the volcanic deposit pending dating. The situation for consideration here likewise matches the “biostratigraphy (PLAN B)” of Copeland , yet replacing the sandstone bed with a volcanic body of interest to be dated.
Furthermore, this work aims to verify whether the interbedded volcanic levels correlate with volcanic events by using a set of criteria to identify primary volcanic deposits in the Lower and Middle Jurassic sedimentary series from the Iberian Chain.
2. Geological setting
2.1. Paleogeographic and geodynamic context
The Pangea supercontinent began to break up in the late Permian, initiating the Alpine Cycle, throughout a series of rifting areas developed with preference along weakness zones such as old suture lineaments between ancient tectonic plates [Quesada and Oliveira 2019]. It continued opening in Jurassic time resulting in two major (Atlantic and Tethyan) tectonic domains. Precise paleogeography of continents and oceans in the western Tethys during the Jurassic is controversial and remains under discussion. The different existing paleogeographic-reconstruction models still demonstrate a certain amount of disagreements, for instance, in the number of oceanic domains (e.g., Ligurian or Alpine Tethys, Betic ocean, Magrhebian Tethys) or the position of fragmented microplates [Poulaki and Stockli 2022, and references therein]. Nevertheless, a consensus about many additional issues has been reached. Thus, about Late Triassic–Early Jurassic times, the Paleotethys had been virtually closed and consumed in favor of the Neotethys Ocean opened to the south. Many terranes (e.g., Cimmerian) drifted northward and then attached to the Euroasian Plate as a result of this Paleotethys Ocean subduction [Schettino and Turco 2011]. The westernmost extension of the Neotethys was the SW–NE trending Alpine Tethys Ocean. This spreading rifted structure constituted a magma-poor rift system [Manatschal et al. 2021] which began to develop during the Early Jurassic [Schettino and Turco 2011].
Iberia, located between the African, Eurasian, and North American plates, played a significant role in shaping plate boundaries in the western Neotethys and Atlantic realms (Figure 1). The western margin of Iberia was constituted by a series of rift basins developed along the future axis of the North Atlantic Ocean [Berra and Angiolini 2014]. The southwestmost segment (Ligurian) of the Alpine Tethys represented the southeast Iberian boundary [Schettino and Turco 2011; Manatschal et al. 2021]. Two transfer zones had fundamental implications for the Iberia-Europe and the Iberia-Africa boundaries: Gibraltar and the North Pyrenean transfer zones [Schettino and Turco 2011; Angrand and Mouthereau 2021]. The Central Atlantic jointed with the westward propagating Ligurian-Alpine mid-ocean ridge around the Pliensbachian [Puga et al. 2011; Schettino and Turco 2011] through the Gibraltar transfer zone. At the Tithonian time, the kinematics of the Gibraltar transfer zone jumped northward to the North Pyrenean transfer zone [Schettino and Turco 2011].
The western half of Iberia was occupied by the emerged Iberian Massif. Its eastern half was constituted by a set of intracratonic basins developed at the same time as the Permo–Triassic rift systems, at the beginning of the Pangea break up, partly reactivating structures inherited from the Variscan orogeny [Osete et al. 2011; Vergés et al. 2019]. In this scenario, the Iberian basin was a carbonate-platform system that constituted the proximal part of the eastern paleomargin of Iberia [Gómez et al. 2004] and recorded alkaline basaltic eruptions interbedded within shallow-marine sediments at 30–35° N latitude [according to Osete et al. 2011] (Figure 1).
The time interval (early Pliensbachian–early Bajocian) in which the studied volcanic deposits were accumulated fits within the first passive margin or post-rifting stage in the Iberian Basin, from the latest Triassic (late Norian) [Sánchez-Moya and Sopeña 2004; Gómez et al. 2019] or Early Jurassic (Sinemurian) [Salas and Casas 1993; Salas et al. 2001] to Middle–Late Jurassic (Callovian–Oxfordian boundary) [Gómez et al. 2019] or latest Oxfordian [Salas and Casas 1993; Salas et al. 2001; Sánchez-Moya and Sopeña 2004]. Most of the total volcanic materials were accumulated in shallow to very shallow marine environments. The volcanic activity in shallow submarine settings tends to be explosive due to the contact between the ambient seawater and the magma and also because the hydrostatic pressure is generally negligible [Cas and Wright 1987; Cas 1992; White et al. 2003; Cas and Giordano 2014]. A part or whole of the volcanic bodies might likely have been redeposited on substrates younger than those in which they were initially deposited (Figure 2).
The Cretaceous and Cenozoic shortening occurred during the subsequent Alpine orogenic stage and gave rise to the Alpine ranges (the Basque-Cantabrian, Pyrenean, Catalonian, Iberian, and Betic ranges). The Iberian Range is a moderately deformed intraplate chain constituted by a NW-oriented folded belt with a low degree of shortening [Gómez et al. 2019]. The Castilian–Valencian Branch and the Aragonese Branch are two NW–SE-oriented elongated areas that form its central and eastern part. The studied area locates at the confluence of both branches (Figure 3A,B,C).
Ancochea et al.  examined the volcanic rocks that crop out on the road connecting the towns of La Puebla de Valverde and Camarena de la Sierra and their surroundings. The lava samples studied are made of olivine basalts and subordinate plagioclase-rich basalts. They mostly show porphyritic textures, displaying varying amounts of olivine and clinopyroxene phenocrysts within a fine-grained microlitic groundmass with plagioclase, clinopyroxene, and olivine. Volcanic rocks are intensively altered as shown by silicification and serpentinization. They are derived from alkaline or middle-alkaline magmas that could have been generated by a moderate degree of melting of the upper mantle reflecting enrichment of incompatible elements. The melt underwent moderate fractionation (11–16%) of olivine and clinopyroxene phases [Ancochea et al. 1988].
Martínez et al. [1996a], Martínez et al. [1996c, d], Martínez et al. [1997a], Martínez et al. [1997b, 1998] and Martínez-González et al.  also studied lava flows and isolated lava bombs included within a pyroclastic pile in the Javalambre ranges area. They are made of porphyritic basaltic rocks with a microlitic groundmass with plagioclase, olivine, and subordinate clinopyroxene minerals. Phenocrysts are made of olivine and titanaugite with concentric zoning. Hematite, ilmenite, titanomagnetite, and spinel are common subordinate opaque minerals. The mineral association indicates an alkaline affinity, confirmed by the chemical composition of clinopyroxenes [Martínez et al. 1996a; Martínez et al. 1996c, d; Martínez et al. 1997a; Martínez et al. 1997b, 1998; Martínez-González et al. 1996]. Ancochea et al. , Martínez et al. [1996a], Martínez et al. [1996c, d], Martínez et al. [1997a], Martínez et al. [1997b, 1998] and Martínez-González et al.  claimed that the volcanism occurred in intraplate domains, being controlled by the development of fracture systems during an extensional period.
2.3. Stratigraphic record
All volcanic bodies are embedded within the Cuevas Labradas (upper part), Barahona, Turmiel, Casinos, and the lower part of the El Pedregal Formations (Figure 3D) [Cortés 2018]. The Cuevas Labradas Formation is constituted by mudstones, bioclastic wackestones to packstones, cross-bedded grainstones, crystalline dolostones, and bindstones (algal mats). Locally, greenish-yellow marls and calcareous breccias are dominant. Overall, facies associations reflect subtidal, restricted lagoon, low- and high-energy intertidal, and supratidal (salt flat) environments. The age interval ranges from the late Sinemurian to the early Pliensbachian (Davoei Zone) up to the late Pliensbachian (Margaritatus Zone) locally [Goy et al. 1976; Gómez et al. 2004].
The Barahona Formation comprises bioclastic wackestones to packstones with chert nodules and hardgrounds, sometimes grainstones, and occasional marly interbeds. Sedimentation took place in relatively shallow subtidal platforms, usually located below the fair-weather wave base but influenced by storm activity and colonized by benthic organisms (mainly crinoids, oysters, and scarce brachiopods). Nevertheless, beach fronts form where bioclastic shoals emerged, showing tractional structures. The Barahona Formation ranges in age from the late Pliensbachian (Margaritatus Zone p.p.) to the early Toarcian (Tenuicostatum Zone p.p.) [Goy et al. 1976; Gómez et al. 2004].
The Turmiel Formation is composed of alternating marls and mudstones to packstones. They were deposited in low-energy, external, and open marine carbonate platforms generally placed below the storm wave base. Lithology, facies associations and diversity of benthic fauna (echinoids, bivalves, bryozoans, brachiopods) indicate widespread deepening of the Iberian carbonate-platform system. The ammonites are more abundant than in the older, previously described formations, especially matching the maximum peaks of accommodation. The Turmiel Formation is Toarcian in age (Tenuicostatum Zone–Bifrons Zone), but the lower part can be late Pliensbachian (Spinatum Zone), and the upper part can locally reach the early Aalenian (Opalinum Zone) [Goy et al. 1976; Gómez and Goy 2000; Gómez et al. 2004].
The Casinos Formation is characterized by mudstones, bioclastic wackestones, and local packstones. In the uppermost part, the formation sometimes shows ferruginous or phosphatic oolites and episodes of regional emersion. Thin marly beds often occur in the lower part of the formation. Facies associations and organisms (crinoids, inoceramids, oysters, brachiopods, gastropods, belemnites, and necroplanktic drifted ammonoid shells) are indicators of external, shallow, and open carbonate platform settings. The Casinos Formation extends from the early Toarcian (Bifrons Zone) to the Aalenian (Murchisonae Zone p.p.) [Gómez et al. 2003, 2004].
The El Pedregal Formation contains mudstones and bioclastic wackestones to packstones, showing locally interlayered marly beds. The base of the formation shows ferruginous or phosphatic oolites. The biotic content includes microfilaments (fragments of thin-shelled bivalves), echinoids (crinoids), gastropods, oysters, brachiopods, sponges, belemnites, and ammonites. The El Pedregal Formation develops in external and shallow carbonate platform environments affected by storms. Its age ranges from the Aalenian (Murchisonae Zone p.p.) to the end of the Bajocian [Gómez and Fernández-López 2004].
2.4. Genetic relationships with surrounding Jurassic volcanism
Similar and roughly coeval volcanism to that of the Iberian Range has been recorded along the southern Jurassic margin of Iberia in what is now the Median Subbetic domain of the Betic Cordillera [García-Yebra et al. 1972; Vera 1988, 2001; García-Hernández et al. 1980; Molina et al. 1998; Molina and Vera 2001; Puga et al. 2004]. Puga et al.  argued that Jurassic basalts and traquibasalts are drawn from fissural volcanism favored by a WSW–ENE trending deep-faulting system.
Volcanic levels are detected from about the Pliensbachian in both the Iberian and Betic basins [Vera et al. 2004; Puga et al. 2011; Gómez et al. 2019; Cortés 2020]. By contrast, they no longer occur from the early Bajocian (late Laeviuscula Zone or Laeviuscula–Propinquans zonal boundary) in the Iberian carbonate-platform system [Cortés 2018, 2021], whereas their presence lasted until the Santonian in the Betic Basin [Molina et al. 1998; Molina and Vera 2008]. Another difference is that lavas, pillow lavas, dikes, and sills were dominant in the Median Subbetic [Gómez et al. 2019] against the pyroclastic explosive deposits that occurred in the Iberian platforms [Cortés 2018]. All the aforementioned suggests the possibility of volcanisms with a common origin in both basins but maybe reservoirs with different capacity.
3. Materials and methods
3.1. Primary and secondary volcanic deposits
There is a need for careful discernment between primary and secondary volcanic deposits to achieve reliable and effective outcomes in the volcanic age estimates. According to White and Houghton  and Sohn and Sohn , primary volcaniclastic deposits are non-reworked ones formed directly from volcanic eruptions (i.e., pyroclastic, autoclastic, hyaloclastic, and peperitic). They contrast with volcaniclastic deposits that are not directly related to eruptions but are reworked, modified, and redeposited by surface or gravitational processes (e.g., tides, waves, currents, or non-eruptive gravitational density flows in the oceanic realm), and are deemed epiclastic or secondary [White and Houghton 2006].
As cementation of volcaniclastic bodies takes time, primary deposits can be eroded, transported and redeposited several times, resulting in submarine volcanic edifices with mixed primary and epiclastic deposits [e.g., Cortés and Gómez 2016, 2018]. Their distinction is often problematic, ambiguous, or even impossible, especially for ancient and weathered sites [Cas and Wright 1987; McPhie et al. 1993; Martínez et al. 1996b; Waitt 2007; Cas and Giordano 2014; Sorrentino et al. 2014; Sohn and Sohn 2019]. Epiclastic layers could be the only relict that provides relevant information about the ancient magmatic activity [Pellenard et al. 1999, 2003; Pellenard and Deconinck 2006; Sorrentino et al. 2014]; hovewer, they could also be redeposited above sediments post-dating the volcanic event itself.
Elementary criteria for distinguishing between primary and secondary volcanic deposits are:
- The presence of continuous lava flows (including pillow lavas) and well-characterized pyroclastic, autoclastic, or hyaloclastic deposits.
- Accretionary lapilli are considered as a distinctive feature of primary volcaniclastic deposits [Cas and Wright 1987; White and Houghton 2006]. They are described in mainly subaerial pyroclastic fall, surge, and flow deposits, formed by accretion of fine ash around a nucleus [Fisher and Schmincke 1984; Cas and Wright 1987]. Indeed, their welded or armored clasts are a diagnostic of subaerial eruptions as pyroclastic flows cannot occur sustainably in aqueous conditions [Fisher and Schmincke 1984; Kokelaar et al. 1984; Cas and Wright 1987; De Goër 2000]. However, it has been suggested that some eruptions of very thick, hot, and dense pyroclastic flows do not mix with ambient water and might keep the heat while being deposited to form subaqueous welded features [Fisher and Schmincke 1984; Kokelaar and Busby 1992]. In any case, accretionary lapilli are unequivocal evidence for syn-volcanic hot deposits of either subaerial or, at best, very shallow subaqueous emplacements [Cas and Wright 1987]. Note that accretionary lapilli can occur as reworked boulders in epiclastic volcanic accumulations far away from the original volcano and primary deposition point [Fisher and Schmincke 1984; Cas and Wright 1987]. To avoid this issue, continuous and non-reworked beds of accretionary lapilli should be considered indicative of primary volcanic deposits.
- Even in submarine emplacements, volcanic accumulations can be lithified before being dismantled and, therefore, they can keep their cone morphology. Although the outcrop conditions are not ideal in ancient examples, observing onlapping wedge-shaped sedimentary units over the volcanic surface indicates a pre-existing primary slope.
- Observing paleo-vent sites linked to volcanic bodies also constitutes a strong argument for primary volcanic accumulation.
- Some volcanic accumulations are not recognizable as volcanic mounds if their deca-kilometric-sized extent greatly exceeds their height (thickness). Large volumes of volcanic materials distributed over great undersea surfaces could indicate primary volcanic subaqueous emplacements. It is hard to understand how primary volcanic deposits may be remobilized and transported far from the origin and later redeposited on younger substrates as huge epiclastic layers.
3.2. Field and laboratory work
The methods used basically consist of fieldwork focused on identifying primary volcanic deposits in order to corroborate the ages of volcanism. To achieve this goal, the geological maps [Cortés 2018] and each of the 13 volcanic levels along the 20 outcrops of the study area were revisited. Volcanic and carbonate rock types, contacts between sediment beds and volcanic bodies, intravolcanic sediment clasts, isolated intravolcanic fossils, and sedimentary structures and ichnological features within intravolcanic sediment beds, were systematically studied and recorded.
Field observations were supplemented with microfacies studies of thin sections obtained both from beds of well-identified lithological units (formations) and intravolcanic sediment clasts. This additional purpose was to determine the genetic provenance of the intravolcanic clasts and check their potential hot contact relations. Thin sections were prepared at the Complutense University of Madrid (UCM) applying standard techniques, such as staining half of each with alizarin red S and potassium ferrocyanide solution. Thin sections were examined under a Nikon Eclipse E400 Pol polarization light microscope and photographed with a coupled Nikon D7100 (24 megapixels) digital microscope camera at the Department of Stratigraphy of the UCM. Petrographic and sedimentological descriptions were carried out according to the Dunham  and Embry and Klovan  classifications.
Taxonomy and taphonomy of carbonate and volcaniclastic internal molds (ammonoids and brachiopods) were performed by Professors Antonio Goy, Sixto Fernández López, Soledad Ureta (UCM), and José Sandoval (Granada University). Most of this paleontologic material was deposited in the Museo Aragonés de Paleontología, Fundación Dinópolis, Teruel (Spain) under the inventory numbers 201/19 and 206/20.
Some figures of this work have been drafted benefiting from the latest available geographic information (topographic and lidar digital elevation—DEM—maps) in the Instituto Geográfico Nacional (IGN) of Spain. The geographic data were processed and managed with ArcGIS software (generation of hillshade maps from DEM with ArcMAP or 3D landscapes with ArcSCENE) and then edited, when necessary, using AutoCAD or Photoshop software programs.
4. Results and interpretations
4.1. Primary volcanic deposits
Lava flows constitute a minority of the total released volcanic products in the studied area. Nevertheless, lava flows have been found in the level V4 (Caudiel [CA] outcrop), level V6 (La Puebla de Valverde.4 [PV.4] outcrop), level V8 (La Puebla de Valverde.2 [PV.2] and La Puebla de Valverde.3 [PV.3] outcrops), and level V12 (Caudiel [CA], Sarrión.1 [SA.1], Sarrión.2 [SA.2], and Sarrión.3 [SA.3] outcrops) (Figure 3C,D). Lavas generally make up thin bodies of a short extent, except for the level V12 in the Caudiel (CA) outcrop and the level V8 in the La Puebla de Valverde.3 (PV.3) outcrop. In the volcanic level V12 of the Caudiel outcrop, a dome-shaped lava flow (about 520 m long in N–S direction and with a maximum thickness of 12 m) was formed (Figure 4A,B). In the volcanic level V8 of the La Puebla de Valverde.3 outcrop, a up to 7 m thick pillow-lavas section has been found (Figure 4C). Well-identified primary pyroclastic deposits (breccia, lapilli, or tuff) have been recognized in all of the studied outcrops (Figure 4D,E), except for the volcanic levels V1, V7, and V10. These latter levels do not show lava flows nor pyroclastic, autoclastic, or hyaloclastic features.
Continuous and non-reworked beds of accretionary lapilli have been found throughout the level V11 (Camarena de la Sierra.1 [CAM.1] outcrop) (Figures 3C,D, and 4F). The chronostratigraphic position of the volcanic level V11 has been linked with an Aalenian (intra-Murchisonae Zone) regional unconformity [Cortés 2018, 2021]. This unconformity separates the Casinos and El Pedregal formations and corresponds to the boundary between the LJ-4 and MJ-1 second-order transgressive-regressive cycles [Fernández-López 1997; Fernández-López and Gómez 2004; Gómez and Fernández-López 2004, 2006]. The shallow marine conditions of that time are compatible with the water-depth conditions going along welding or above the water-air interface eruptions.
Well-characterized volcanic mounds, as well as volcanic flanks onlapped by younger carbonate beds, have been found in the level V11 (Camarena de la Sierra.3 [CAM.3] outcrop), level V12 (Caudiel [CA], Pina-Barracas.1 [PI-BA.1], Sarrión.1 [SA.1], Sarrión.2 [SA.2], and Sarrión.3 [SA.3] outcrops), and level V13 (Abejuela [AB] and Llíria [LLÍR] outcrops) (Figures 3C,D, and 5).
At least two possible near-vent sites have been identified both in the volcanic level V9 (Camarena de la Sierra.5 [CAM.5] outcrop) and in the level V12 (Caudiel [CA] outcrop). Some features, such as: (i) local breakage able to form subrounded clasts from unconsolidated carbonate beds, probably as a response to the explosive eruption, and (ii) injection of irregular lenses of volcanic matter into the broken sediments, are visible in the volcanic level V9 from the Camarena de la Sierra.5 (CAM.5) outcrop (Figures 3C,D, and 6A,B,C). Moreover, centimeter- to meter-thick angular calcareous blocks embedded into lava bodies as a result of conduit wall rock fragmentation, partly assimilated and exhibiting peperitic and fluidification textures, are observed in the volcanic level V12 from the Caudiel (CA) outcrop (Figure 6D,E). They would correspond to the composite clasts of White and Houghton , formed by the interrelationship between magma and sediment (fragments of peperite).
The most evident large volcanic extents are observed within the volcanic levels V4, V5, V6, V9, and V11 from the La Puebla de Valverde (PV) and Camarena de la Sierra (CAM) outcrops (Figure 3C,D). For example, the volcanic level V4 in the CAM.1 outcrop shows 2100 m of visible linear length along NE–SW, the volcanic levels V5 and V6 in the PV.4 outcrop extend over an area of about 12 km2, the volcanic level V9 in the CAM.5 outcrop covers an area of around 5 km2, and the level V11 in the CAM.1 outcrop displays 4500 m of visible linear length in a NE–SW direction. Other large volumes of volcanic deposits, although across more reduced areas, can also be related to the V2 (Camarena de la Sierra.3 [CAM.3], Camarena de la Sierra.4 [CAM.4], and Camarena de la Sierra.5 [CAM.5] outcrops) and V3 volcanic levels (Pina-Barracas.2 [PI-BA.2] and Pina-Barracas.3 [PI-BA.3] outcrops). For example, just over 2000 m of visible linear length along NE–SW for the level V2 in the CAM.5 outcrop, and over 660 m from east to west for the level V3 in the PI-BA.2 outcrop.
4.2. Features indicative of secondary volcaniclastic deposits?
The occurrence of some elements, such as sediment clasts, beds, or fossils within the volcaniclastic piles, could be quoted to argue their secondary epiclastic origin. Whether clasts, fossils, and beds are indicative of secondary volcanic deposits or not is evaluated in the following sub-sections and accompanying figures. Interpretations are based on the results of sedimentological field observations and microfacies analyses.
4.2.1. Fossils with volcaniclastic infill
The presence of fossils could be used to postulate reworking in the volcaniclastic deposits. However, it is undeniable that fossils can occur in primary volcaniclastic piles, even in solid basalts [e.g., Nayudu 1971]. The fossilization happened within the volcanic deposits, and their durability as preserved fossils is estimated as very low but not impossible (Figure 7A,B). The fossil preservation on the superficial levels of the volcanic mass would merely denote fall, shallow burial, and an internal volcanic filling of the shells, but not necessarily reworking processes. However, fossils, especially nektonic ones, found to a certain depth indicate a higher draft in reworking processes. Fossils with volcaniclastic infill will always be more recent than the volcanic body where they are included. Nevertheless, fossil gathering that occurred shortly after the volcanic accumulation may be considered contemporary for most chronostratigraphic dating purposes.
4.2.2. Sediment clasts
Sediment clasts included in the volcaniclastic bodies are generally assumed to be fragments of country rocks enclosed and ejected along with the pyroclasts. However, it cannot be ruled out the possibility that primary volcaniclastic deposits and pre-existing sediments have been reworked and redeposited together, resulting in a mixture of volcaniclasts and sedimentary clasts. Thus, their presence within volcaniclastics could support that such volcanic bodies are reworked epiclastic deposits.
Carbonate clasts of sedimentary origin, ranging from a few centimeters to several decimeters in size, are commonly observed all over the volcaniclastic levels from the Pliensbachian to the Bajocian in the studied area. They show high morphologic variability (from subrounded or ellipsoidal to subangular) and often have a recrystallized external appearance along with a frequent concentric black alteration halo.
One of the most exceptional sites concerning the frequency and size variation of intra-volcanic sedimentary clasts is the Abejuela outcrop (volcanic level V13) (Figure 3C,D). Several typical rock types from the pre-volcanic lithostratigraphic units were identified as clasts, as shown by their lithologic features (Figure 7C) and their fossil content (e.g., Brasilia sp. along with Malladaites sp., Prisnorhynchia rabesaxensis, and Pseudogibbirhynchia mutans, from the Aalenian) (Figure 7D,E,F).
Thin sections observations indicate that clasts are clearly made of the pre-volcanic units (sometimes much older than the chronostratigraphic position of the volcanic deposit) (Figures 8A–F, 9A–D). They further revealed peperitic textures commonly related to magma intrusions and mingling with the still unconsolidated and wet host sediments. It demonstrates that clasts and magmas were in initial contact (Figure 9E,F).
These observations consistently indicate that the sedimentary clasts included in the volcanic bodies are former wall rocks of volcanic conduits remobilized by explosive volcanic processes. That is, the presence of sediment clasts does not have to be indicative of secondary epiclastic deposits, but rather the opposite.
4.2.3. Sediment beds
The presence of centimeter- to meter-thick sedimentary beds included in the volcaniclastic piles requires different consideration than sediment clasts. Detailed observations from these commonly single beds have revealed the widespread occurrence of hummocky cross-stratification (HCS), sedimentary structure typically considered diagnostic of storm deposits in the shallow marine realm [e.g., Dott and Bourgeois 1982; Hunter and Clifton 1982; Duke 1985; Haines 1988; Monaco 1992; Sami and Desrochers 1992; Dumas and Arnott 2006]. These storm deposits are clearly related to sedimentary processes. Tempestites surrounding volcaniclastic mounds contain clasts of volcanic origin but may also contain fossils. The latter are interpreted to be younger than the volcanic event except if they were reworked from older strata [e.g., reelaborated ammonites of Fernández-López 1984].
Relationships between primary and secondary volcaniclastic deposits are sometimes found within a single pile: a cm-thick calcareous bed located at the lower and distal part of a volcanic pile can be seen in Figure 10A (Caudiel outcrop). It is made of a bioclastic (mainly bivalves, ammonoids, and crinoids) and intraclastic wackestone–packstone (Figure 10B) interpreted as a tempestite. Intraclasts are volcanic epiclasts. It is overlain by a few meters of volcaniclastic deposits of the same facies as the lower volcanics.
In the La Puebla de Valverde.4 outcrop (Figure 10C), a primary pyroclastic deposit (lapilli) is observed at the lower part of the section (Figure 10D). Above, a cm-thick bed with abundant epiclasts is interpreted as another storm deposit (Figure 10E), covered by fine-grained volcaniclastic deposits with faint traction structures (Figure 10F).
Therefore, since the storm beds (both in Caudiel and La Puebla de Valverde.4 outcrops) are sedimentary in nature, the volcaniclastic material accumulated above them must have been remobilized and redeposited later, as their features and texture often demonstrate, although they are not always easy to distinguish [Cas and Wright 1987; McPhie et al. 1993; Martínez et al. 1996b; Waitt 2007; Cas and Giordano 2014; Sorrentino et al. 2014; Sohn and Sohn 2019]. The volcaniclastic deposits postdating the storm beds can be interpreted as sliding from the summit of the mounds triggered by the wave activity or collapse mechanisms. In any case, the volcanic section predating the storm beds should be deemed as a primary deposit. Instead, the stretches overlying the sedimentary storm layers are non-primary reworked (epiclastic) products.
In short, it has been established that: (a) both the occurrence of sediment clasts and fossils with volcaniclastic infill is not necessarily indicative of secondary, reworked, and epiclastic deposits, and (b) a sediment layer interbedded within the flank of a volcaniclastic pile clearly marks the boundary between primary (below it) and epiclastic (above it) volcanics.
5.1. Stratigraphic ages of volcanic deposits and timing of volcanism
The analysis of volcanic deposits according to diagnostic criteria for primary volcanic deposits is summarized in Figure 11. As can be seen, most of the 13 volcanic levels must be witnesses of primary volcanism. Almost all of them meet one or several criteria in one or more outcrops. The exceptions are the volcanic levels V1, V7, and V10 because they do not meet any of the above criteria. These three volcanic levels show in all of the outcrops and across their entire thickness a mixture of volcanic and non-volcanic components (thin mud-carbonate drapes separating uneven or lenticular volcanic debris, Figure 12A, thin lenticular beds of mudstone with plane parallel lamination, Figure 12B), tractional sedimentary structures (undulate stratification with bioclastic and sedimentary particles, Figure 12C, cross and planar lamination, Figure 12D), complete invertebrate marine fossils embedded in the volcano-sedimentary groundmass, and hallmarks of bioturbation.
They are portions derived from volcanic edifices either preserved in the subsurface or already entirely eroded, with the doubt as to whether they are coeval or more recent than the volcanism age.
The volcanic level V7 occurs across the Camarena de la Sierra.5 (CAM.5) outcrop, reaching only a centimetric thickness where it crops out. The same thicknesses are observed for volcanic level V10, present in the La Puebla de Valverde.4 (PV.4) and Sarrión.3 (SA.3) outcrops. It could be argued that V10 is reworked from the volcanic level V9 in the La Puebla de Valverde.4 (PV.4) outcrop but cannot be demonstrated for the Sarrión.3 (SA.3) outcrop where V9 is not observed. Finally, the volcanic level V1 has only been locally recorded in the Camarena de la Sierra.4 (CAM.4) outcrop. If its volcano-sedimentary origin is assumed, the occurrence of an older volcanic phase should be investigated but is not confirmed by outcrops.
Since the youngest volcanic level (V13) verifies several criteria for a primary volcanic deposit, the resulting period of volcanism spans the early Pliensbachian (or a bit older) to the early Bajocian. The number of syn-eruptive volcanic phases would correspond with the number of interbedded volcanic deposits (i.e., 13) or slightly lower (possibly between 10 and 12, if one, two, or all three levels in question—V1, V7, V10—do not represent actual volcanic events).
5.2. Spatio-temporal sequence of volcanic emissions along structural trends
The Jurassic volcanic activity (early Pliensbachian–early Bajocian) is supposed to develop in a post-rift or passive margin stage [Salas et al. 2001; Gómez et al. 2019] and therefore in a tectonic quiescence period with decreasing subsidence rates. However, Gómez , Fernández-López and Gómez , Gómez et al. , Gómez and Goy , Gómez and Fernández-López , and Gómez et al.  noticed the presence of two systems of NW–SE and NE–SW trending faults active during the Early and Middle Jurassic through the central and southern Iberian Range, which delimited depocenters and topographic highs.
An additional rift pulse associated with magmatism is referred by Van Wees et al. , who claimed that the number of rifting pulses of low magnitude is much higher than was usually estimated. A period of rapid tectonic subsidence during the Pliensbachian–Toarcian interval would be followed by decreasing subsidence rates from the Aalenian to the Oxfordian. Although neglected in other studies, much of the magmatic early Pliensbachian–early Bajocian time interval falls within it.
Regional mapping [Gautier 1974; Abril et al. 1975, 1978; Campos et al. 1977; Lazuen and Roldán 1977; Adrover et al. 1983] demonstrates that the volcanic rocks are emplaced along three of the mentioned linear structural discontinuities, two NW–SE trending faults [Caudiel and Alcublas fault zones, Gómez 1979] and one oriented NE–SW [Teruel Fault Zone, Fernández-López and Gómez 2004] (Figure 3C). It might be speculated that the NW–SE oriented Caudiel and Alcublas fault zones could be laterally connected to the Ligurian mid-ocean ridge axis by transform faults. By contrast, the NE–SW (or NNE–SSW) Teruel Fault Zone runs roughly orthogonal to the Jurassic extensional direction.
Based on the biostratigraphic dating carried out by Cortés [2018, 2020, 2021], Figure 13 shows the spatial distribution of the subaqueous emission centers that change along time. The volcanism mostly focused on the Teruel Fault Zone during the Early and earliest Middle Jurassic, being sporadically recorded towards the Caudiel Fault Zone, such as the volcanic episodes V3, V4, and V11 (Figure 13).
This spatial distribution changed during the Middle Jurassic. The active volcanism migrated southeastward from the Teruel Fault Zone to the Caudiel Fault Zone during the emplacement of the volcanic episode V12 (Aalenian–Bajocian, Concavum–Discites, boundary in age).
Then, the volcanism migrated southwards from the Caudiel Fault Zone to the Alcublas Fault Zone. It is not until the uppermost Laeviuscula Zone or the Laeviuscula–Propinquans zonal boundary that the Alcublas Fault Zone underwent magmatic activity with the emission of the volcanic episode V13, which appears solely arranged along this fault zone (Figure 13).
5.3. Geodynamic context of the Jurassic magmatism within the Iberian Range
Intraplate magmatism is, in principle, hard to explain within the frame of plate tectonics [Farnetani and Hofmann 2011; Lee and Grand 2012]. There is general agreement that intraplate volcanism is triggered by decompression melting and related to deep mantle processes—plumes (hotspots)—, although perhaps not as deep as initially thought [Foulger and Natland 2003], or off-axis magmatism.
The volcanism recorded in the Iberian Range meets a number of requests on the specifics of hotspots and off-axis magmatism, such as intraplate setting, basalts enriched in incompatible elements (e.g., Perfit and Davidson ), or magmatic crustal penetration controlled by crustal structures such as fault systems [e.g., Mutschler et al. 1998].
Typical off-axis magmatism develops over oceanic crust on the faulted flanks of mid-ocean ridges [e.g., Sohn and Sims 2005; Canales et al. 2012; Toomey 2012; Carbotte et al. 2012, 2016; Choi et al. 2021]. Off-axis magmatism has been reported as occurring up to 20 km beyond the ridge axis in the East Pacific Rise [Turner et al. 2011], up to 40 km in the Gulf of Aden [Guillard et al. 2021], or by more than 200 km in the volcanic fields of the Kamar-Daban, the Udokan and the Vitim Plateau [Yang et al. 2018].
Longer distances than those mentioned could undoubtedly separate the studied area from the Ligurian mid-ocean ridge. Moreover, the studied area locates over the rifted continental crust. This suggests that the Jurassic volcanism of the Iberian Range might be originated from hotspot magmatism (Figure 1). The Iberian magmatic region, located at the junction between the Betic and Ligurian basins, constituted a zone of diffuse continental extension that recorded distributed phases of rift until Late Cretaceous times [Angrand et al. 2020]. This scenario, also supported by Van Wees et al.  for the Pliensbachian–Toarcian interval, could have encouraged the rise and extrusion of magma. The lack of magmatism during the second Mesozoic Iberian rifting stage—from the Callovian–Oxfordian boundary [Gómez et al. 2019] or the latest Oxfordian [Salas and Casas 1993; Salas et al. 2001; Sánchez-Moya and Sopeña 2004] to the Albian—could be tentatively explained by magmatic chamber depletion. However, it is somewhat paradoxical that magmatism was present in the Iberian Basin during a generally agreed inter-rifting stage.
Using biochronostratigraphic methods applied in sediments hosting volcanic bodies to date volcanic events is effective when there is a good record of fossils with chronostratigraphic value, such as ammonites. The results can be applied to other problematic dating cases or to double-check for existing dating. It is essential to keep in mind that some volcanic levels might not match exactly with any volcanic event but that they would have been eroded from the prime accumulation and redeposited on younger substrates than those in which they were primarily stored.
The implementation of field criteria as tools for the identification of primary volcanic accumulations has made it possible to establish a reliable time interval for the Jurassic volcanism (ca. 20 Ma) in the southeastern Iberian Range, at zone or even subzone ammonite scale: from the early Pliensbachian (Jamesoni Zone or later) to the early Bajocian (uppermost Laeviuscula Zone or Laeviuscula–Propinquans zonal boundary). It unlocked the possibility to discuss the spatio-temporal migration of active volcanism at the scale of the study area.
Through precise dating of volcanic episodes, it can be observed that the timing of magmatic eruptions was not simultaneous in the three Fault Zones affecting the area. Exceptionally, the emission of volcanic products becomes synchronous in two of them. Furthermore, what has been noticed is that the volcanic outcrops are shown grouped around preferred locations that change over time. Firstly, it shifted towards the southeast from the Teruel Fault Zone to the Caudiel Fault Zone around the Aalenian–Bajocian boundary and then southward (or towards the SSW) from the Caudiel Fault Zone to the Alcublas Fault Zone around the early Bajocian (uppermost Laeviuscula Zone or Laeviuscula–Propinquans zonal boundary).
Conflicts of interest
The author has no conflict of interest to declare.
Thanks to Antonio Goy Goy, Sixto Fernández López, and José Sandoval Gabarrón for the taxonomic determination of the brachiopod and ammonite specimens referred in this work. The author would like to thank the useful comments, suggestions, and revisions of Emmanuel Masini and two anonymous reviewers, which contributed to a much improved paper. The careful editorial handling of Michel Campillo is greatly appreciated.