Discussion of the palaeoecology of belemnites, or the validity of the isotopic data obtained from belemnite calcite for the calculation of palaeotemperatures do not fall within the scope of this research, but the use of belemnite calcite as a proxy is generally accepted and widely used as a reliable tool for palaeothermometry in most of the Mesozoic. However, belemnite palaeoecology constitutes a source of conflicts because, due to the fact that these organisms are extinct, there is a complete lack of understanding of fossil belemnite ecology (Rexfort and Mutterlose, 2009). Belemnites lived as active predators within swimming life habitats. Nevertheless, several authors (Anderson et al., 1994; Mitchell, 2005; Wierzbowski and Joachimiski, 2007) have proposed a bottom-dwelling lifestyle on the basis of oxygen isotope thermometry, similar to modern sepiids which show a nektobenthic mode of life. This is contradicted by the occurrence of various belemnite genera in black shales which lack any benthic or nektobenthic organisms due to the existence of anoxic bottom waters (i.e. the Lower Jurassic Posidonienschiefer, see Rexfort and Mutterlose, 2009), a fact that indicates that belemnites presented a nektonic mode of life rather than a nektobenthic (Mutterlose et al., 2010). As Rexfort and Mutterlose (2009) stated, it is unclear whether isotopic data from belemnites reflect a surface or a deeper water signal, and we are unaware whether the belemnites mode of life changed during ontogeny. Similarly, Li et al. (2012) concluded that belemnites were mobile and experienced a range of environmental conditions during growth; furthermore, these authors stated that some belemnite species inhabited environmental niches that remain unchanged, while other species had a more cosmopolitan lifestyle inhabiting wider environments. To complete the scenario, Mutterlose et al. (2010) suggested different lifestyles (nektonic versus nektobenthic) of belemnite genera as indicated by different shaped guards. Short, thick guards could indicate nektobentic lifestyle, elongated forms fast swimmers, and extremely flattened guards a benthic lifestyle.

The study by Ullmann et al. (2014) hypothesizes that belemnites (Passaloteuthis) of the Lower Toarcian Tenuicostatum Zone had a nektobenthic lifestyle and once became extinct (as many organisms in the Early Toarcian mass extinction) were substituted by belemnites of the genus Acrocoelites supposedly with a nektonic lifestyle, which these authors attribute to anoxia.

The isotopic studies performed on present-day cuttlefish (Sepia sp.), which are assumed to constitute the group most equivalent to belemnites, reveals that all the specimens (through their δ18O signal) perfectly reflect the temperature characteristics of their habitat (Rexfort and Mutterlose, 2009). Also the studies of Bettencourt and Guerra (1999), performed in cuttlebone of Sepia officinalis, conclude that the δ18O-obtained temperature agreed with changes in seawater temperature, thus supporting the use of belemnites as excellent tools for calculation of palaeotemperatures.

It seems that at least some belemnites could swim through the water column, reflecting average temperature and not necessarily only bottom or surface water temperatures. In any case, rather than single specific values, in the present paper comparisons of average temperatures to define the different episodes of temperature changes are used.

The Upper Sinemurian, Pliensbachian and Lower Toarcian deposits of the Rodiles section comprise couplets of bioclastic lime mudstone to wackestone limestone and marls. These limestones occasionally contain bioclastic packstone facies concentrated in rills. Limestones, generally recrystallized to microsparite, are commonly well stratified in beds whose continuity can be followed at the outcrop scale, as well as in outcrops several kilometres apart. However, nodular limestone layers, discontinuous at the outcrop scale, are also present. The base of some carbonates can be slightly erosive, and they are commonly bioturbated, to reach the homogenization stage. Ichnofossils, especially Thalassinoides, Chondrites and Phymatoderma, are also present. Marls, with CaCO3 content generally lower than 20 % (Bádenas et al., 2009, 2012), are frequently grey coloured, occasionally light grey due to the higher proportion of carbonates, with interbedded black intervals. Locally brown coloured sediments are present, more often in the Upper Sinemurian.

The ammonite-based biochronostratigraphy of these deposits in Asturias was performed by Suárez-Vega (1974), and the uppermost Pliensbachian and Toarcian ammonites by Gómez et al. (2008) and by Goy et al. (2010a, b). Preliminary biochronostratigraphy of the Late Sinemurian and the Pliensbachian in some sections of the Asturian Basin has been reported by Comas-Rengifo and Goy (2010), and herein we summarize the result of over 10 years of bed by bed sampling of ammonites in the Rodiles section, which provided a precise time constraint for the climatic events described in this work.

The ammonites collected enabled recognition of all the standard Late Sinemurian, Pliensbachian and Early Toarcian chronozones and subchronozones defined by Elmi et al. (1997) and Page (2003) for Europe. The section is generally expanded and ammonites are sufficiently common to constrain the boundaries of the biochronostratigraphical units. Exceptions are the Taylori-Polymorphus subchronozones that could not be separated, and the Capricornus-Figulinum subchronozones of the Davoei Chronozone, partly due to the relatively condensed character of this Chronozone. Most of the recorded species belong to the NW Europe province but some representatives of the Tethysian Realm are also present.

Belemnites in the Rodiles section generally show an excellent degree of preservation (Fig. 4) and none of the prepared samples were rejected, as only the non-luminescent parts of the belemnite rostrum not affected by diagenesis were selected. It has been assumed that the biogenic calcite retains the primary isotopic composition of the seawater and that the belemnite migration, skeletal growth, sampling bias, and vital effects are not the main factors responsible for the variations obtained.

The cross-plot of the δ18O against the δ13C values (Fig. 5) reveals a cluster-type distribution, showing a negative correlation coefficient (-0.2) and very low covariance (R2=0.04), supporting the lack of digenetic overprints in the diagenetically screened belemnite calcite analysed.

The carbon isotopes curve reflects several oscillations throughout the section studied (Fig. 3). A positive δ13Cbel shift, showing average values of 1.6 ‰ is recorded from the Late Sinemurian Densinodulum to part of the Macdonnelli subchronozones (from metre 0 to 21 in Fig. 3). From the latest Sinemurian Aplanatum Subchronozone (Raricostatum Chronozone) up to the Early Pliensbachian Valdani Subchronozone of the Ibex Chronozone, average δ13Cbel values are -0.1 ‰, delineating an approximately 1–1.5 ‰ relatively well marked negative excursion (from meter 21 to 57 in Fig. 3). In the late Ibex and in the Davoei chronozones, the δ13Cbel curve records background values of around 1 ‰, with a positive excursion at the latest Ibex Chronozone and the earliest Davoei Chronozone (from meter 62 to 67 in Fig. 3).

At the Late Pliensbachian the δ13Cbel values tend to outline a slightly positive excursion (from meter 71 to 97 in Fig. 3), interrupted by a small negative peak in the latest Spinatum Chronozone (from meter 98 to 103 in Fig. 3). The Early Toarcian curve reflects the presence of a positive δ13Cbel trend which develops above the stratigraphical levels represented herein, up to the Middle Toarcian Bifrons Chronozone (Gómez et al., 2008) and a negative excursion recorded in bulk carbonates samples.

The δ18Obel values show the presence of several excursions throughout the Late Sinemurian to the Early Toarcian (Fig. 3). From the Late Sinemurian to the earliest Pliensbachian interval, a negative excursion of around 1 ‰, showing values generally below -1 ‰(from meter 0 to 18 in Fig. 3) with peak values up to -3 ‰ has been recorded in Sinemurian samples located immediately below the stratigraphic column represented in Fig. 3. In most of the Early Pliensbachian Jamesoni and the earliest part of the Ibex chronozones, δ18Obel values are quite stable, around -1 ‰ (from meter 18 to 56 in Fig. 3), but another negative excursion of approximately 1–1.5 ‰, with peak values up to -1.9 ‰, develops along most of the Early Pliensbachian Ibex and Davoei chronozones, extending up to the base of the Late Pliensbachian Margaritatus Chronozone (from meter 56 to 76 in Fig. 3). Most of the Late Pliensbachian and the earliest Toarcian are characterized by the presence of a significant change. In this interval (from meter 76 to 103 in Fig. 3) a positive excursion in the order of 1.5 ‰ δ18Obel, with frequent values of around 0 ‰, and positive values up to 0.7 ‰, were assayed. The oxygen isotopes recorded a new change in its tendency in the Early Toarcian, where a prominent δ18Obel negative excursion, about 1.5–2 ‰ with values up to -3 ‰, has been verified (from meter 103 to 110 in Fig. 3).

The detailed biostratigraphical analysis, based on the succession of the Pliensbachian ammonoids assemblages allowed construction of a scale of reference that has facilitated the location of the different palaeoclimatic events recognized in the present research.

The five biochronozones of the standard scale constituting the Pliensbachian of the Subboreal/NW Europe Province (Dommergues et al., 1997; Page, 2003) have been recognized in the Rodiles section. For the first time, these biochronozones have been subdivided into 14 subchronozones whose boundaries have been corrected in many cases with respect to previous studies. In most cases these boundaries have now been established with a low margin of uncertainty.

With regard to previous research (Suárez-Vega, 1974; Comas-Rengifo and Goy, 2010) the Taylori and Brevispina subchronozones of the Early Pliensbachian have been characterized in this study for the first time, and the boundary between the Valdani and the Luridum subchronozones, usually difficult to distinguish in the Asturian Basin, has been clearly recognized. In the Late Pliesbachian, where the record of Amaltheidae is quite complete, the subchronozone Apyrenum of the Spinatum Chronozone has been characterized and the boundary between the Subnodosus and Gibbosus subchronozones has been precisely established.

The δ13Cbel carbon isotope excursions (CIEs) found in the Asturian Basin, can be followed in other sections across Western Europe (Fig. 6). The Late Sinemurian positive CIE was also recorded in the Cleveland Basin of the UK by Korte and Hesselbo (2011) and in the δ13Corg data of the Wessex Basin of southern UK by Jenkyns and Weedon (2013).

The Early Pliensbachian δ13Cbel negative excursion extending from the Raricostatum Chronozone of the latest Sinemurian to the Early Pliensbachian Jamesoni and part of the Ibex chronozones (Fig. 6), correlates with the lower part of the δ13Cbel negative excursion reported by Armendáriz et al. (2012) in another section of the Asturian Basin. Similarly, the δ13Cbel curve obtained by Quesada et al. (2005) in the neighbouring Basque–Cantabrian Basin shows the presence of a negative CIE in a similar stratigraphical position. In the Cleveland Basin in the UK, the studies on the Sinemurian–Pliensbachian deposits conducted by Hesselbo et al. (2000), Jenkyns et al. (2002) and Korte and Hesselbo (2011) reflect the presence of this Early Pliensbachian δ13Cbel decrease in values. In the Peniche section of the Lusitanian Basin of Portugal, this negative CIE was also recorded by Suan et al. (2010) in brachiopod calcite, and in bulk carbonates in Italy (Woodfine et al., 2008; Francheschi et al., 2014). The magnitude of approximately 1.5–2 ‰ of this negative excursion appears to be quite consistent across the different European localities.

Korte and Hesselbo (2011) pointed out that the Early Pliensbachian δ13C negative excursion seems to be global in character, resulting from the injection of isotopically light carbon from some remote source, such as methane from clathrates, wetlands, or thermal decomposition, thermal metamorphism or decomposition of older organic-rich deposits. However, none of these possibilities has yet been documented.

Higher in the section, the δ13C values are relatively uniform, except for a thin interval, around the Early Pliensbachian Ibex–Davoei zonal boundary, where a small positive excursion (the Ibex–Davoei positive excursion, previously mentioned by Rosales et al., 2001 and by Jenkyns et al., 2002) can be observed in most of the δ13C curves summarized in Fig. 6, as well as in the carbonates of the Portuguese Lusitanian Basin (Silva et al., 2011).

The next CIE involves a positive excursion of around 1.5–2 ‰, well recorded in all the correlated Upper Pliensbachian sections (the Late Pliensbachian positive excursion in Fig. 6) and in bulk carbonates of the Lusitanian Basin (Silva et al., 2011; Silva and Duarte, 2015) and in the Apennines of Central Italy (Moretinni et al., 2002). This CIE also partly coincides with the δ13Corg reported by Caruthers et al. (2014) in western North America. Around the Pliensbachian–Toarcian boundary, a negative δ13C peak is once again recorded (Fig. 6). This narrow excursion was described by Hesselbo et al. (2007) in bulk rock samples in Portugal, and tested by Suan et al. (2010) in the same basin and extended to the Yorkshire (UK) by Littler et al. (2010) and by Korte and Hesselbo (2011). If this perturbation of the carbon cycle is global, as Korte and Hesselbo (2011) pointed out, it could correspond with the negative δ13C peak recorded in the upper part of the Spinatum Chronozone in the Asturian Basin (present paper); with the negative δ13C peak reported by Quesada et al. (2005) in the same stratigraphical position in the Basque–Cantabrian Basin, and with the δ13C negative peak reported by van de Schootbrugge et al. (2010) and Harazim et al. (2013) in the French Grand Causses Basin.

Finally, the Early Toarcian is characterized by a prominent δ13C positive excursion that has been detected in all the sections considered herein, as well as in some South American (Al-Suwaidi et al., 2010) and Northern African (Bodin et al., 2010) sections. This positive CIE is interrupted by a negative excursion of approximately 1 ‰ δ13Cbulk located around the Tenuicostatum-Serpentinum zonal boundary.

The origin of the positive excursion has been interpreted by some authors as the response of water masses to excess and rapid burial of large amounts of organic carbon rich in 12C, which led to enrichment in 13C of the sediments (Jenkyns and Clayton, 1997; Schouten et al., 2000). Other authors ascribe the origin of this positive excursion to the removal from the oceans of large amounts of isotopically light carbon as organic matter into black shales or methane hydrates, resulting from ebullition of isotopically heavy CO2, generated by methanogenesis of organic-rich sediments (McArthur et al., 2000).

Although δ13C positive excursions are difficult to account for (Payne and Kump, 2007), it seems that this positive CIE cannot necessarily be the consequence of the widespread preservation of organic-rich facies under anoxic waters, as no anoxic facies are present in the Spanish Lower Toarcian sections (Gómez and Goy, 2011). Modelling of the CIEs performed by Kump and Arthur (1999) shows that δ13C positive excursions can also be due to an increase in the rate of phosphate or phosphate and inorganic carbon delivery to the ocean, and that large positive excursions in the isotopic composition of the ocean can also result from an increase in the proportion of carbonate weathering relative to organic carbon and silicate weathering. Other authors argue that an increase of δ13C in bulk organic carbon may reflect a massive expansion of marine archaea bacteria that do not isotopically discriminate in the type of carbon they use, giving rise to positive δ13C shifts (Kidder and Worsley, 2010).

The origin of the Early Toarcian δ13C negative excursion has been explained by several papers as resulting from the massive release of large amounts of isotopically light CH4 from the thermal dissociation of gas hydrates. Hesselbo et al. (2000, 2007), Cohen et al. (2004) and Kemp et al. (2005), associated it with the massive release of gas methane linked with the intrusion of the Karoo-Ferrar large igneous province onto coalfields, as proposed by McElwain et al. (2005) or with the contact metamorphism by dykes and sills related to the Karoo-Ferrar igneous activity into organic-rich sediments (Svensen et al., 2007).

Martinez and Dera (2015) proposed the presence of fluctuations in the carbon cycle during the Jurassic and Early Cretaceous, resulting from a cyclicity of ∼ 9 My linked to a great eccentricity cycle, amplified by cumulative sequestration of organic matter. Nevertheless, this ∼ 9 My cycle has not been evidenced in the Pliensbachian deposits of several parts of the World (Ikeda and Tada, 2013, 2014) and cannot be evidenced in the Pliensbachian deposits of the Asturian Basin either. The disruption of this cyclicity recorded during the Pliensbachian could be linked to chaotic behaviour in the solar system (Martinez and Dera, 2015) possibly due to the chaotic transition in the Earth–Mars resonance (Ikeda and Tada, 2013). Data from Japan suggest that this disruption, which developed from the Hettangian to the Pliensbachian (Ikeda and Tada, 2013, 2014) was possibly linked to the massive injection of CO2 from the eruptions of the Central Atlantic Magmatic Province to the Karoo-Ferrar eruptions (Prokoph et al., 2013) which destabilized the carbon fluxes, reducing or dephasing the orbital imprint in the δ13C over millions of years (Martinez and Dera, 2015).

Seawater palaeotemperature calculation from the δ18O values obtained reveals the occurrence of several isotopic events corresponding to relevant climatic oscillations across the latest Sinemurian, the Pliensbachian and the Early Toarcian (Fig. 7). Some of these climatic changes could be of global extent. In terms of seawater palaeotemperature, five intervals can be distinguished. The earliest interval of these corresponds to a warming period developed from the Late Sinemurian up to the earliest Pliensbachian. Most of the Early Pliensbachian is represented by a period of “normal” temperature, close to the average palaeotemperatures of the interval studied. A new warming period is recorded in the Early–Late Pliensbachian transition, and the Late Pliensbachian is represented by an important cooling interval. Finally the Early Toarcian coincides with a severe (super)warming interval, linked to the important Early Toarcian mass extinction (Gómez and Arias, 2010; García Joral et al., 2011; Gómez and Goy, 2011; Fraguas et al., 2012; Clémence, 2014; Clémence et al., 2015; Baeza-Carratalá et al., 2015). The average palaeotemperature of the latest Sinemurian, Pliensbachian (palaeolatitude of 32∘ N) and Early Toarcian (palaeolatitude of 40∘ N), calculated from the δ18O values obtained from belemnite calcite in the present study, is 15.6 ∘C. As mentioned above, some belemnites could swim through the water column, and the palaeotemperatures calculated do not necessarily correspond only with the temperatures of the bottom or surface waters, but also the average temperature.