The field campaign at Playa de los Caballos (Figure 1C) was conducted in November 2017 and Eighty-three (83) samples were collected at the nominal Cuchía Section, starting from the uppermost beds of the Umbrera Fm to the middle portion of the Patrocinio Fm, including the entire black shale unit previously reported to correspond to OAE1a (Wilmsen, 2005; Najarro et al., 2011b; García-Mondéjar et al., 2015a; García-Mondéjar et al., 2015b). The sampling resolution is ∼0.82 m with slightly closer intervals that target more indurated interbeds. Because of active mass wasting along the outcrop, extra care was given to the sampling process by removing ∼5—8 cm of the exposed surface to avoid contamination. Colors assigned to the rocks on dry samples follow the standard “Rock-color chart” of the Geological Society of America (Goddard et al., 1984). The studied portion of the section includes essentially ∼2.36 m of limestone layers (medium greenish gray 5GY 5/1) at the base, and the remainder overlying 65.64 m is in sharp contrast consisting of dark shales (greenish black 5GY 2/1) with interbeds of calcareous shales and marlstones (dark greenish gray 5GY 4/1).

The total carbon (TC), total inorganic carbon (TIC), and total organic carbon (TOC) content were measured for all 83 samples at Florida International University (FIU). The analysis was conducted using a LECO CR-412 carbon analyzer following the procedures laid out by Sanchez-Hernandez and Maurrasse (2014). A pure calcite standard reference material (C64-500, Fisher Scientific) was used to calibrate the instrument by producing a best-fit calibration curve. In preparation for the analysis, all samples were powdered using a Bell-Art micromill with a tungsten-carbide chamber and blade. TC content was acquired by introducing approximately 0.3 g of powdered sample to the carbon analyzer. For TIC, before processing the samples with the carbon analyzer, ∼0.3 g of powder for each sample was first placed in a muffle furnace at 560 °C for 2 h to remove the organic carbon content. The results produced by the instrument are reported as CaCO₃ wt%. Therefore, to obtain TOC as C wt%, the following formula was used: T O C C w t % = T C − T I C × 8.33 , where the 8.33 comes from the molecular weight of CaCO₃ divided by the molecular weight of C (100.086/12.011) (Ponton Guerrero, 2006). Here, TIC values are used to assign a lithologic classification based on the nomenclature after Sanchez-Hernandez and Maurrasse (2014), where: limestone, >65% CaCO₃; marly-limestone, 60%–65% CaCO₃; marlstone, 30%–60% CaCO₃; calcareous shale, 10%–30% CaCO₃; and shale, 0%–10% CaCO₃.

The stable carbon isotope on the organic fraction (δ¹³C_org) was conducted at the Stable Isotope Laboratory (SIL) at FIU using a Thermo Delta C EA-IRMS following the procedures after Swart et al. (2019). Prior to the analysis, powdered samples were decarbonated by acidification with 10% HCl until all the carbonate was removed, then rinsed with deionized (DI) water to raise the pH. The reproducibility of the δ¹³C_org values is ±0.1‰ based on glycine standards. The results for the δ¹³C_org are given in per mil (‰) relative to the Vienna Peedee Belemnite standard.

Elemental analyses were carried out on thirty (30) samples targeting the various lithologies throughout the studied interval of the studied outcrop. Major (Al, Ti, Si), biolimiting (Fe, P), and redox sensitive trace elements (RSTEs; V, Cr, Ni, Mo, U) were analyzed at FIU’s Trace Evidence Analysis Facility (TEAF) following the methodology developed by Arroyo et al. (2009). In preparation for the analysis 1.000 g of powdered material were spiked with two internal standard solutions: 300 µl indium metal (1,000 ppm in 3% HNO₃) and 350 µl scandium oxide (1,000 ppm in 3% HNO₃) and set to dry on a heating block at 80 °C for 48 h. Next, samples were further powdered and homogenized using a tungsten carbide ball-mill for 30 min and pressed into pellets ready for analysis. Concentrations were determined by Laser Ablation-Inductively Coupled Mass Spectrometry (LA-ICP-MS) using the ELAN DRCII (Perkin Elmer LAS) with quadrupole in standard operating mode. An LSX-500 with a compact Nd:YAG laser (266 nm) operating with a 200 μm spot size at 10 Hz was used for the ablations. Standard reference materials PACS-2, a marine sediment (National Research Council of Canada, Ottawa, Canada); and two soil reference materials: SRM NIST2710 (Montana Soil) and NIST2704 (Buffalo River Sediment) were used for calibration and to check for accuracy. Results were averaged for each sample based on five ablations relative to their standard deviations using the GLITTER software. The overall average precision for the analysis was <5% (1.7%–14%) relative standard deviation, with a 0.2% average error based on replicate standards. Major and bioessential elements are reported in mg/g and RSTEs are reported as parts per million (ppm).

The clay and mineralogical contents were measured for nine samples (n = 9) using X-ray diffraction (XRD) at the Illinois State Geological Survey, Geochemistry Section, Prairie Research Institute, University of Illinois. The clays were analyzed on the <2 µm fraction to determine the variations of smectite, illite, kaolinite, and chlorite to the basin. The analyses were performed using a Scintag^® XDS2000 diffractometer. Sample preparations followed the procedures laid out by Moore and Reynolds (1989). Step-scanned data were collected from 2° to 34° 2θ with a fixed rate of 2° per min with a step size of 0.05°2θ determined from oriented and glycolated clay slides. For the bulk minerals, the samples were prepared in the same manner described by Socorro and Maurrasse (2022), where the samples first were micronized for 10 min in a McCrone micronized mill with deionized water. After drying they were homogenized and packed into end-loading sample holders and then were analyzed on a Scintag XDS2000 diffractometer with data collection set from 2° to 60° 2 θ with a fixed time of 5 s per 0.05° 2θ for each sample. The mineral phases identified include quartz, potassium feldspar (K-feldspar), plagioclase feldspar (P-feldspar), calcite, dolomite, siderite, pyrite/marcasite, and bulk clay. All resulting traces were quantified using the semi-quantitative data reduction software from Materials Data Inc. (MDI) known as Jade+ and reported as relative percent.

Twenty (20) samples targeting different lithologies were selected for biomarker analysis. The preparation for the biomarker analysis follows the procedures proposed by Socorro and Maurrasse (2019) and further explained in Herdocia and Maurrasse (2022). The total lipids were extracted by adding 5 g of powdered sample to a 10 mL solution of dichloromethane (DCM) and methanol (9:1) and sonicated for 30 min. The solution was then centrifuged, and the lipids were collected. This process was repeated twice for each sample. Elemental sulfur was removed by adding copper chippings to the extracts. Finally, the lipids were concentrated by evaporating the DCM and methanol, and the concentrate was left in 1 mL of n-hexane solution. The biomarker analysis was carried out at FIU’s Advanced Mass Spectroscopy Facility using an Aligent 6,890 gas chromatography (GC) equipped with an HP-5MS UI capillary column (30 m, 0.250 mm, 0.25 μm). Detection was performed by an Aligent 5,973 single-quadrupole mass spectrometer (MS), using electron impact in positive mode. The MS was operated in Full Scan Mode between m/z 40 and 400. 5μL of each sample were injected into the GC with an injection temperature of 280 °C. The temperature program used for the biomarker analysis was as follows: hold at 65 °C for 2 min, increase to 300 °C at a rate of 4°C/min, and hold at 300 °C for 8 min.

As applied at previous sites in the Iberian Peninsula (Socorro et al., 2017; Socorro and Maurrasse, 2019) we also used normal alkane (n-alkane) chain lengths as a proxy to better characterize the relative proportion of land-derived in contrast with marine-derived OM. The ratio of terrestrial to aquatic extracts or TAR is calculated by dividing the sum of the longer odd-numbered n-alkanes by the sum of the shorter chain odd-numbered n-alkanes (Bourbonniere and Meyers, 1996; Peters et al., 2005; Holtvoeth et al., 2010). T A R = n C 27 + n C 29 + n C 31 / n C 15 + n C 17 + n C 19

Values below 1 indicate a predominance of marine derived OM, while values above 1 indicate a predominance of terrestrial OM.

The relative abundance of odd to even numbered carbons for the long chain n-alkanes (≥nC₂₄) was used to determine the carbon preference index (CPI) (Bray and Evans, 1961): C P I = [ n C 25 + n C 27 + n C 29 + n C 31 + n C 33 n C 24 + n C 26 + n C 28 + n C 30 + n C 32 + n C 25 + n C 27 + n C 29 + n C 31 + n C 33 n C 26 + n C 28 + n C 30 + n C 32 + n C 34 ] / 2

The odd-to-even predominance of the long chain n-alkanes (>nC₂₄) was determined using the OEP (2) after Peters et al. (2007) which was modified after Scalan and Smith’s (1970) OEP equation as follows: O E P 2 = n C 25 + 6 n C 27 + n C 29 / 4 n C 26 + 4 n C 28

All correlation come from Microsoft Excel 2016, using Pearson’s r equation (1896): r = ∑ x i − x ¯ y i − y ¯ ∑ x i − x ¯ 2 ∑ y i − y ¯ 2

In the formula, x_i and y_i are the corresponding sample values, and their corresponding sample averages are indicated by x ¯ and y ¯ . All r-values are in the range of −1 to 1. A value of −1 is equivalent to a perfect negative correlation, 0 indicates that there is no correlation, and an r-value of 1 is equals a perfect correlation.

TIC, TOC, δ¹³C_org were measured in all 83 samples; their values are plotted in Figure 3 (Table 1). TIC content for the 67.27 m studied section range from 0.1% to 86.0% (Figure 3B). TIC in the lower 6 m of the section fluctuates from 5.5% in the marlstone beds to 86.1% in the limestones and includes the Umbrera Formation (0–2.36 m) as well as the lowermost Patrocinio Fm. The overlying 38.7 m (SC-17-13 to SC-17-55; 7.04 m–44.72 m) is a part of the lower Patrocinio Formation which yielded TIC values between 3.4% and 25.4%. TIC in the upper Patrocinio Formation (44.72 m–67.27 m) is more variable with fluctuations from 68.1% down to 0.1%. A closer look in the Upper Patrocinio shows that above 44.72 m up to 55.02 m (SC-17-56 to SC-17-69), TIC values fluctuate between 8.3% and 52.4%. These high values are followed by a drop in TIC between 1.1% and 3.9% up to the height of 56.69 m (SC-17-70 to SC-17-74). Sequentially to 63.62 m (SC-17-75 to SC-17-79) TIC increases again from 15.9% up to 68.1%. The top of the studied section, above 56.69 m (SC-17-80 to SC-17-83) shows a significant drop in TIC values between 0.1% and 0.9%.

TOC values vary between 0.3% and 1.2%, with an average of 0.6% (Figure 3C). The lowest part of section (0–16.68 m; SC-17-1 to SC-17-24) shows increasing TOC values between 0.4% and 0.9%, whereas in the succeeding ∼4.56 m (17 m–21.56 m; SC-17-25 to SC-17-32), TOC values drop to a low fluctuation pattern (0.5%–0.6%). The subsequent ∼16.51 m (22.16 m–38.67 m; SC-17-33 to SC -17-49) includes more pronounced variations in TOC values, between 0.6% and 1.2% and contains two peaks in TOC, 1.2% and 1.1% at 28.56 m and 38.67 m, respectively. In the overlying ∼5.58 m (39.59 m–45.17 m; SC-17-50 to SC-17-55) TOC values are relatively constant (0.7%–0.8%). The uppermost ∼21.27 m of the studied section up to 67.27 m (SC-17-56 to SC-17-83) shows TOC variations with high frequencies, between 0.3% and 1.0% with an overall upward increasing trend.

As shown in Figure 3D, carbon isotopic values (δ¹³C_org) vary between −26.6‰ and −22.8‰, with an average of −25.8‰. The δ¹³C_org values fluctuate from −25.2‰ (0 m) to a small positive value of −24.7‰ at 2.36 m (Umbrera Fm), which is followed by an abrupt negative shift to −26.2‰ at 2.61 m, at the base of the Patrocinio Fm. This negative inflection extends from 2.61 m to 21.6 m with an average of −26.4‰. The following isotopic values show an abrupt positive shift of 1‰ from −26.3‰ at 22.16 m to −25.3‰ at 24.61 m. This positive shift continues with a slow decline up to 61.23 m, where the value is −25.7‰. Another strong positive shift of 2.9‰ occurs from 61.23 m reaching up to −22.8‰ at 67.27 m.

Results for the analysis of major elements and RSTEs are plotted in Figure 4 and Figure 5 (Table 2). Commonly elemental concentration for major, bioessential, and RSTEs are measured to assess terrigenous supplies, nutrient availability for productivity and the influence of reducing conditions during sedimentation.

The major elements fluctuate with lithology as the highest values coincide with the shales while the lowest correspond with the limestones, and they share a strong negative correlation with TIC (r = Al, −0.95; Si, −0.99; Ti, −0.97). Al concentrations range between 3.06 mg/g to 88.01 mg/g and average 56.86 mg/g (Figure 4C). Si ranges between 19.46 mg/g and 205.92 mg/g (Figure 4D) and correlates strongly with Al (r = 0.95). Ti concentrations correlate strongly with Al (r = 0.96) and fluctuate between 0.20 mg/g and 3.86 mg/g (Figure 4E). These terrestrially derived elements show a positive correlation with TOC (r = Al, 0.51; Si, 0.56; Ti, 0.52) as shown in Figure 4.

Of the biolimiting elements (Fe, P), Iron follows a pattern like Al (r = 0.71) and shows a positive correlation with TOC (r = 0.52) (Figure 4G). In general, Fe is higher in the shales and decreases in the limestone, as indicated by its negative correlation with TIC (r = −0.65). By contrast, phosphorus has no apparent correlation with either Al (r = −0.21) or TOC (r = −0.01) and ranges from 0.17 to 0.89 mg/g with an average of 0.32 mg/g (Figure 4F).

The temporal distribution of the RSTEs (V, Cr, Ni, Mo, and U) is shown in Figure 5. Vanadium, and U show trends compatible with Al (V, r = 0.94 and U, r = 0.76) and of the measured RSTEs they share the strongest correlation with TOC (V, r = 0.50 and U, r = 0.51). Molybdenum behaves independently of the lithology, TOC, and the major elements.

Figure 6 shows the results for the mineralogical analysis carried out for 9 samples at stratigraphic heights: 5.06 m (SC-17-10), 7.04 m (SC-17-13), 14.21 m (SC-17-21), 28.56 m (SC-17-41), 36.79 m (SC-17-47), 47.84 m (SC-17-60), 57.27 m (SC-17-71), 61.23 m (SC-17-76), and 66.36 m (SC-17-82). The bulk of the clay mineralogy is composed of Illite (64.9%–78.6%; Figure 6B). Smectite and illite contents seem to drop at stratigraphic heights with high TOC but increase where TIC is high, whereas Kaolinite and chlorite percentages decrease where TIC is higher. Chlorite shows minor variations between 6.9% and 10.9%. Kaolinite content fluctuates from 11.3% to 21.5%. Smectite is in the lowest abundance of the clays (0.8%–3.8%). Figure 6C shows the bulk mineralogical abundances averaged for the 9 analyzed samples. The most abundance mineral is quartz (42%) followed by clay minerals (17%) then siderite (14%), calcite (10%), P-feldspar (6%), K-feldspar (5%), dolomite (4%), and pyrite/marcasite (2%).

The twenty samples analyzed for n-alkanes and the acyclic isoprenoids pristane (Pr) and phytane (Ph) (Figure 7). Twelve (12) samples at the respective heights of 0.6 m, 3.85 m, 7.04 m, 12.07 m, 15.04 m, 19.05 m, 24.61 m, 38.67 m, 50.28 m, 53.64 m, 57.28 m, and 67.27 m (Figure 7C), have a bimodal distribution with one peak centered at nC₁₅ and the other between nC₂₅ and nC_28, (Figures 7D,G–I). The remaining 8 samples (heights: 28.56 m, 33.57 m, 43.05 m, 46.0 m, 47.84 m, 49.02 m, 61.23 m, and 63.62 m) display a right skewed (positive) unimodal distribution (Figures 7E,F). The chromatograms for four representative examples (Figure 8) reveal mostly low unresolved complex mixture (UCM) for the aliphatic fraction. The TAR values throughout the section range from 0.06 up to 1.41 with an average of 0.43 (Table 3; Figure 7C). In the lower part TAR values vary from 0.48 up to 1.41 at stratigraphic height 19.05 m, which corresponds to the highest values attained. The TAR values decrease from 1.41 to 0.15 at 25.56 m and remain below 0.5 until 63.62 m. Generally, the n-alkane distribution has an inverse relationship with TIC (r = −0.46) as higher TIC coincides with lower TAR values, whereas low TIC corresponds with higher TAR values.

The n-alkanes with chain lengths greater than nC₂₄ have an average CPI of 0.82 (from 0.6 to 1.4) and an average OEP (2) of 0.75 (from 0.5 to 1.5) indicating that the succession has mostly an even carbon number preference (Table 3). Pristane and Phytane ratios (Pr/Ph) are listed on table 3 and range from 1.4 to 5.0 with an average of 2.19. Values for Pr/nC₁₇ and Ph/nC₁₈ are low, averaging 0.31 and 0.23, respectively.

Temporal CIEs (δ¹³C_org and δ¹³C_carb) coeval with OM-rich intervals recorded in Aptian sediments worldwide are widely accepted to represent global disruptions in the carbon cycle (e.g., Scholle and Arthur, 1980; Arthur et al., 1985; Jenkyns, 1995; Kuhnt et al., 1998; Menegatti et al., 1998; Leckie et al., 2002; de Gea et al., 2003; Godet et al., 2006; Michalík et al., 2008; Tejada et al., 2009; Sanchez-Hernandez and Maurrasse, 2016; Hu et al., 2020; Socorro and Maurrasse, 2022). Potential local variations in the amplitude of the carbon isotope curve have also been examined because of modulating factors such as terrestrial OM and species–specific fractionation (Ravelo and Fairbanks, 1995; Taylor et al., 2001; Anderson et al., 2005), OM source and biochemical pathways (e.g., Pardue et al., 1976; Cranwell, 1981; Ponton et al., 2014; Feakins et al., 2007; Feakins et al., 2016; Basu et al., 2015) as well as diagenesis. While these previous factors may affect the Cuchía site, another factor related to tectonic controls on the long-term carbon isotope mass balance (Shields and Mills, 2017) can be ruled out, because covariations of the defined carbon isotopic segment associated with OM-rich sediments of the studied section have been satisfactorily used on a wide geographic scale (Millán et al., 2009; Moosavizadeh et al., 2014; Sanchez-Hernandez and Maurrasse, 2014; Sanchez-Hernandez and Maurrasse, 2016; Socorro and Maurrasse, 2019). Since the isotopic segments have also been well-tested and calibrated to magnetostratigraphy and biostratigraphy (Channell and Erba, 1992; Channell et al., 2000; Herrle et al., 2004; Godet et al., 2006), the effectiveness of the δ¹³C_org and the δ¹³C_carb curve can be considered a robust and useful tool for cross correlation between sites. Thus, despite the modulating factor from terrestrial sources the carbon isotopic curve is a reliable chronostratigraphic proxy for the Aptian stage. Based on this premise, we use these segments to contrast the studied site with other well-studied sections, namely, the El Pui section (Sanchez-Hernandez and Maurrasse, 2016), the Cismon section, and the Roter Sattel section (Menegatti et al., 1998) to confirm the Cuchía timeline in the global context (Figure 10).

The δ¹³C_org values at the base of the studied section fluctuate from −25.2‰ (0 m) to a value of −24.7‰ at 2.36 m at the top of the Umbrera Fm which can be assigned to the uppermost part of C2 (Figure 10D). This spike is followed by an abrupt negative shift to −26.2‰ at 2.61 m and reaches the most negative value of ∼-26.5‰ at 21.56 m. A similar pattern is observed at Roter Sattel (Switzerland), toward the end of C2 (∼3.5 m) with a C-isotope value of ∼-27.2‰, then at 4 m it reaches a peak of −26.4‰, marking the end of segment C2, which is followed by a negative excursion to a low of −28.8‰ indicative of the end of C3. In the Cismon section (Italy), segment C2 ends with a maximum of ∼-24.0‰ which is succeeded by fluctuating values that end with a sharp negative value of ∼-28.4‰ marking the close of C3 (Menegatti et al., 1998). By comparison, in the El Pui section (Catalunya, Spain), segment C2 ends with a maximum of −25.0‰ at 163 m, followed by a sharp decline to ∼-27.2‰ corresponding with the start of the negative trend characteristic of C3 which reaches its lowest value of ∼-28.0‰ at 199 m (Sanchez-Hernandez and Maurrasse, 2016).

The remarkable distinction of the C3 negative trend recorded in the lower part of the Patrocinio Fm (2.61m - 21.56 m) is the low amplitude fluctuation of the δ¹³C_org values from ∼-26.6‰ to ∼-26.1‰, a shift of only ∼-1.8‰. By comparison, the three coeval sites yield the following changes: Roter Sattel ∼-2.4‰; Cismon ∼-4.4‰; El Pui ∼-3.0‰ (Figures 10A–C). As we will discuss further the cause of this discrepancy, sediments of the NBCB at the Cuchía site recorded the least negative values of this distinct CIE that is the hallmark of the C-isotope curve marking the onset of OAE1a.

A similar divergence is recorded in the Cuchía section for the conventional negative CIE assigned to C4 which is marked by a 1‰ positive shift from −26.3‰ at 22.16 m to −25.3‰ at 24.61 m. By contrast, in the Roter Sattel section segment C4 (Figure 2 in Menegatti et al., 1998) is defined by ∼3.4‰ positive inflection from −28.8‰ to about −25.4‰. Similarly, in the Cismon section (Figure 3 in Menegatti et al., 1998) segment C4 is represented by a ∼2.5‰ positive shift from ∼−28.4‰ to ∼−25.9‰, whereas in the El Pui section segment C4 is characterized by ∼3.7% positive shift extending ∼13 m, (199 m–∼212 m) with an isotope value ranging from −27.9‰ to −24.2‰ at 212 m (Sanchez-Hernandez and Maurrasse, 2016).

The succeeding δ¹³C_org values assigned to segment C5 in the Cuchía section start with ∼-25.3‰ at the end of segment C4 and remain steady around −25.7‰ ending with a value of ∼-25.7‰ at 61.72 m (Figure 10D). By comparison, segment C5 in the Roter Sattel section yields isotope values between −25.6‰ and −24.8‰ with little variation (∼0.8‰), whereas in the Cismon section isotope values for segment C5 fluctuate between ∼−26.0‰ and −23.6‰ (∼2.4‰) with no sudden increase or decrease within the segment. In the conjoint El Pui - El Pujal section (Sanchez-Hernandez and Maurrasse, 2016; Socorro and Maurrasse, 2019; Socorro and Maurrasse, 2022) segment C5 varies between −25.3‰ and ∼-22.6‰, with an amplitude of ∼2.7‰. The final 5.5 m portion of the studied Cuchía section may be attributed to carbon-isotope segment C6 because it shows another strong 2.9‰ positive shift with values from −25.7‰ up to −22.8‰ at 67.27 m (Figure 10D). Likewise, in the Roter Sattel section a sharp positive CIE of ∼2.0‰ defines carbon-isotope segment C6 with an increase from about ∼-25.6‰ to ∼-23.6‰ (Figure 10A). A comparable abrupt fluctuation is also recorded in the Cismon section which shows a ∼3.2‰ increase from −25.2‰ to ∼-22.0‰ (Figure 10B).

In essence, the pattern of δ¹³C_org fluctuation in the studied Cuchía section correlated to Roter Sattel, Cismon and El Pui permits to identify segments C2 to C6 (Figure 10) coinciding with previous studies (Najarro et al., 2011a; Najarro et al., 2011b). Carbon isotope values at the Cuchía site vary from −26.6‰ to −22.8‰, giving a spread of 3.8‰, thus revealing that the most negative values occur within segment C3, identical to the record in the other correlative sections. Non-etheless, the values are less negative in the NBCB compared to elsewhere (Figure 10). The main difference is the overall amplitude of the isotopic values in C3 that is more positive at Cuchía. The Cismon, Roter Sattel, and the El Pui sections share very similar values, the lowest value in all three sites is about −28‰ within segment C3 and their maximum values in C6 are around −22‰, giving all three sites about a six‰ spread, or nearly 40% larger. In the following section we further discuss the possible factors that may have caused the differences between the range of δ¹³C_org values in Cuchía and the correlative sites.

The true OM signal may be affected by thermal alteration that brings bias to the data; therefore, the degree of thermal maturity is an important parameter to assess the reliability of the OM present. Commonly, bimodal distributions in the n-alkanes are lost with increasing degree of thermal maturity because hydrocarbons are selectively removed (Peters et al., 2005). Given that 12 out of 20 of the chromatograms for the biomarkers (Figure 7C) have a bimodal distribution, therefore the degree of thermal degradation is likely moderate in the studied section.

The degree of thermal degradation may be further tested by calculating the relative abundance of odd to even numbered carbons in the long chain n-alkanes using the CPI and OEP (2) (Peters et al., 2005). In general, as CPI and OEP (2) values approaches ∼1 thermally maturity increases, while values significantly above 1.0 are indicative of an odd preference, and values significantly below 1.0 point to an even carbon preference (Bray and Evans., 1961; Peters et al., 2005). The long chain n-alkanes show a CPI that ranges from 0.61 to 1.55 with an average of 0.89, and OEP (2) values from 0.47 to 1.47 with an average of 0.75 (Table 3). Both ratios suggest that the samples have an even carbon number preference and corroborates some degree of thermal alteration as the values are close to 1.0. Usually, OM derived from terrigenous sources display a predominance of odd-over-even carbon number of long chain n-alkanes (≥nC₂₅), which is disclosed particularly for nC₂₇, nC₂₉, and nC₃₁, but this preference may disappear during early diagenetic processes (Johnson and Calder, 1973; Nishimura and Baker, 1986; Kuhnt et al., 1998; Peters et al., 2005). In the Cuchía section most of the long chain n-alkanes show an even-over-odd preference, except at a height of 12.07 m and 67.27 m. The lack of odd-over-even carbon preferences throughout the section points to low to moderate degree of diagenetic alteration that may not significantly alter their original characteristics.

By comparison, biomarkers from three other sites elsewhere within the Basque-Cantabrian Basin, southeast and east of Cuchía (Quijano et al., 2012), similarly showed that C₂₂-C₃₀ n-alkanes had an even carbon number preference (OEP = 0.69—0.80; Chaler et al., 2005). This even carbon preference was attributed to strong reducing conditions which can favor the formation of even-carbon-number from fatty acids and waxy alcohols. Yet, as discussed above, RSTEs in the Cuchía section correlate strongly with the major elements, thus indicating that this part of the basin may not have experienced strong reducing conditions. The mechanism proposed by Chaler et al. (2005) probably played a role in the even carbon preference at Cuchía but was also modulated by thermal alteration.

In other parts of the NBCB, previous biomarker results based on hopane and sterane ratios at the Puentenansa section (named La Florida section in Najarro et al., 2011b) to the southwest of the Cuchía section (Figure 1A), at a level equivalent to the Patrocinio Fm, contained thermally mature OM within the peak oil generation window (Quijano et al., 2012). Given that the NBCB experienced uneven subsidence rates (Najarro et al., 2011b) it can be inferred that different parts of the NBCB experienced dissimilar thermal histories since basin subsidence leads to increased depth of burial and thus higher geothermal effects (Feinstein, 1981). In fact, as shown by Najarro et al. (2011a), Najarro et al. (2011b) the Puentenansa section experienced slightly higher subsidence rates than the Cuchía area. Thus, based on the bimodal distribution, the CPI and the OEP (2) values, we can surmise that the rocks in the Cuchía section are probably either within the oil window, or perhaps in the early oil window and support the accuracy of the isotopic data.

To elucidate the intricate paleoenvironmental evolution of the NBCB recorded in the Cuchía section it is important to consider the possible modulating effects of paralic influences as suggested in previous studies (Najarro et al., 2011b). We use n-alkanes as a proxy to characterize the type of OM preserved and determine the temporal contribution of the land derived components to the basin. The chain length of the n-alkanes are well recognized proxies that can be used to determine the source of the OM because short chain n-alkanes (≤nC₁₉) are mainly derived from marine plankton and/or microbial communities (Cranwell, 1973; Cranwell et al., 1987; Forster et al., 2004; Peters et al., 2005), whereas chain lengths between nC₂₀ to nC₂₅ tend to be produced by marine and non-marine aquatic macrophytes (Ficken et al., 2000), and terrestrial plants produce longer chain n-alkanes (≥nC₂₅) (Peters et al., 2005).

The chronological distribution of the biomarker data from the Cuchía section revealed a range of n-alkanes from nC₁₁ to nC₃₃ (Figure 7), thus implying a mixed OM content throughout the length of the studied interval. Of the 20 samples analyzed, 8 have a unimodal distribution centering on nC₁₅ indicative of a predominance of OM derived from marine organisms. Twelve (12) samples have a bimodal distribution with one peak centered at nC₁₅ and the second peak centered between nC₂₅ and nC₂₈ (Figure 7C). In all cases, the short-chain n-alkanes are most abundant, therefore they are indicative of a predominance of OM mostly derived from microbial phytoplankton and aquatic macrophytes, although the long-chain n-alkanes also disclose significant contribution of terrestrial plants (Dañobeitia et al., 1990; Frakes et al., 1992; Kuhnt et al., 1998; de Gea et al., 2003; Forster et al., 2004). This mix composition is corroborated by the Pr/nC₁₇ and Ph/nC₁₈ plot (Figure 9) which shows that the OM is mostly within marine type II kerogen, as well as mixed type II/III kerogen emblematic of both marine and terrestrial contribution (Peters et al., 2005). Sample SC17-28 (19.05 m) is the only one that plots in terrigenous type III kerogen. Nevertheless, uncertainty may arise in the interpretation because these ratios can be affected through biodegradation since bacteria preferentially degrade hydrocarbons. Typically, the n-alkanes will be removed first, especially the low molecular weight n-alkanes, followed by the acyclic isoprenoids (Pr and Ph) (Peters et al., 2005). The extent of such changes can be easily identified in the chromatograms by the UCM which becomes increasingly more prominent as these compounds are removed. The samples at Cuchía have very low UCM humps at all analyzed levels as well as higher proportions of n-alkanes to the acyclic isoprenoids (Figure 8), therefore they were preserved with minimal biodegradation.

Further assessment of the provenance relative to the contribution of terrigenous versus marine OM can be more distinctly discerned using the TAR (Figure 7C) which compares the relative abundances of short chain n-alkanes (nC₁₅, nC₁₇, nC₁₉) derived from marine OM (Cranwell, 1973; Cranwell et al., 1987; Socorro and Maurrasse, 2019; Socorro and Maurrasse, 2020), to the long chain n-alkanes (nC₂₇, nC₂₉, nC₃₁) produced by terrestrial plants (Peters et al., 2005). Higher TAR values correspond to increased terrigenous OM (Bourbonniere and Meyers, 1996). In the Cuchía samples TAR generally follows lithology (TAR vs. TIC, r = −0.46), thus indicating heightened terrigenous OM input at levels with decreased carbonate content. Additionally, the TAR temporal variation follows a trend comparable to the Al content (r = 0.53), revealing that terrestrial pulses are the main delivery method for the terrestrial biomarkers (Hedges et al., 1997). The highest TAR values tend to correspond primarily with the lowermost part of the sequence (below 24.61 m) with a ratio up to 1.41. Although the average TAR value of 0.43 reveals that the OM was predominantly produced by marine microbial phytoplankton, long chain n-alkanes (≥nC₂₅) are found at all measured levels, which corroborates sustained input from terrigenous sources, especially below 24.61 m where TAR values are highest. This data agrees with the Pr/nC₁₇ and Ph/nC₁₈ plot (Figure 9) as well as the gymnosperm pollens revealed in a previous study (Najarro et al., 2011b).

In the Cuchía sediments the C-isotopic values correlated with the sections elsewhere are comparable in all segments except within segment C3 where they are somewhat less negative, which begs the question concerning the possible causes for the higher isotopic values observed in this study. Given that the environment of deposition of the Cuchía sequence is delta influenced, we postulate that the variations observed arise from the averages of combined effects of modulated global factors due to the diverse OM from a complex ecosystem that developed in the basin.

A cursory review of the mechanism of carbon metabolic absorption will help to better understand the complexity of the carbon fractionation process and how it may have affected the residual OM preserved in the Cuchía sediments. It is generally accepted that δ¹³C_org values depend on the different pathways of C-sequestration which are critical to the ratio obtained as an endmember of the photosynthetic process. In this relation, varying kinetic fractionation of δ¹³C_org develops during photosynthesis. Basically, when terrestrial plants utilize atmospheric CO₂, they preferentially absorb the lighter ¹²C which is easier to metabolize than ¹³C because of their relative differential masses. Fractionation is further influenced by the photosynthetic pathway of carbon as plants may utilize the Calvin Cycle and they are designated C₃ plants because the CO₂ is initially fixed into a compound that contains three carbon atoms (phosphoglyceric acid) before entering the Calvin cycle. Terrestrial C₃ plants have δ¹³C_org compositions more depleted in ¹³C, yielding lower values of −22‰ to −33‰, generally averaging between −26‰ and −28.5‰ (Bender, 1971; Boutton, 1991; and others); they are interpreted as more negative, or with lower isotopic values. Climatic factors can also affect the carbon isotope values, as for instance, dominant C₃ plants in the present Gangetic plain (India) yield δ¹³C_org values from −27‰ to −31.4‰, (x̄=−28.9‰ ± 1.1‰), which decrease with increasing mean annual precipitation (Basu et al., 2015). Certain plants may also use the Hatch-Slack pathway where the CO₂ forms a simple 4-carbon organic compound (Oxaloacetic acid), classified as C₄ plants; this fractionation leads to higher δ¹³C values in the −10‰ to −20‰ range (Bender, 1971). However, Bowes et al. (2002) hypothesized that aquatic C₄ photosynthesis possibly arose before its onset in terrestrial plants, and fossil plants indicate that C₄ plants originated much later, in the Paleogene Period, ∼34—24 Ma (Sage, 2004). Therefore, contribution from C₄ plants can be ruled out as an enhancement factor for the higher δ¹³C_org isotopic values in the Cuchía carbon isotopic archive. Thus, the enhancement factor of the carbon isotopic values in the Cuchía organic record must lie within the OM supplied by terrestrial C₃ photosynthesizers.

In this context, the remaining issue to be clarified in the Cuchía deposit is the photosynthetic mechanism that may have generated more positive δ¹³C_org values from OM of C₃ photosynthesizers. The C₃ metabolic pathway with direct atmospheric-sourced CO₂, aquatic plants and marine algae acquire carbon from dissolved inorganic carbon (DIC), usually in the form of dissolved CO₂ or HCO₃ ⁻ (White, 2013). In this case, ¹³CO₂ dissolves more readily than ¹²CO₂ yielding a fractionation of about +0.9‰. Further fractionation of +7‰ to +8‰ occurs during hydration and dissociation of H₂CO₃, which also enriches ¹³C in the DIC (White, 2013). Since the Cuchía section is a marine deposit, the OM is most likely derived predominately from pelagic in situ phyto- and zooplankton. But, the long chain n-alkanes and elevated TAR values (Figure 7) indicate that it was also strongly influenced by a delta (Rat, 1988), thus implying that riverine fluxes contributed to land-sourced OM which tend to have lower isotopic values. However, contrary to what would be expected, the carbon isotope negative excursion during C3 in Cuchía is not as pronounced as it is elsewhere (Figure 10). This anomaly can be explained because previous palynologic study of the section reveals the presence of gymnosperms (Najarro et al., 2011b). Indeed, modern aquatic analogs have shown that the carbon isotopic content for pond cypress trees (Taxodium ascendens) appears to respond to changes in precipitation, and more positive δ¹³C_org values correlated with higher precipitation, yielding carbon isotopic values between −21.77‰ and −20.65‰ (Anderson et al., 2005). It is well demonstrated that the time associated with OAE1a was characterized by increased precipitation (e.g., Barron et al., 1981; Barron et al., 1995; Sanchez-Hernandez and Maurrasse, 2016 and others), accordingly we can surmise that such climatic conditions may have also affected the metabolic processes of the Aptian gymnosperms as observed in the present, which may explain the cause of the discrepancy in the record showing higher carbon isotopic values in the sedimentary archive of the Cuchía section.

It has been well documented that the OAE interval is characterized by increased weathering rates and terrestrial fluxes due to enhanced precipitation, but how did the global greenhouse climate affect the supply of the terrigenous component in the expression of OAE1a in the delta influenced NBCB? This issue is addressed by measuring the temporal variability of the major elements (Al, Si, Ti), the clay and bulk mineral distributions, biophile elements (P, Fe), and the RSTEs because they can be used to infer periods of enhanced precipitation as their main delivery mechanism as well as productivity and the redox conditions at the time of deposition.