Thermal demagnetization results indicate the presence of two paleomagnetic components. A low-T component unblocks typically below 200°C and has shallow inclinations. This component is interpreted as an unstable component acquired during drilling, and will not be discussed further. A characteristic remanent magnetization (ChRM) directed toward the origin of the orthogonal vector component plot with both positive and negative directions is identified above 200°C and up to 600°C, which suggests that magnetite is the main carrier of the NRM (Figure 5). Stable ChRM directions with unblocking temperatures up to 600°C are typically associated with larger NRM intensities (Figures 5A,B). When the NRM is weaker, the ChRM has maximum unblocking temperatures of <450°C but still with a well-defined linear trend directed to the origin of the demagnetization plots (Figures 5C–E). Regardless, samples with reliable ChRM directions are labeled as Type 1. Type 2 samples have less clearly identified ChRM directions (e.g., Figure 5F), and are not considered further. Type 1 ChRM directions have both positive and negative inclinations regardless of NRM intensity, which suggests that the ChRM provides a reliable record of geomagnetic polarity reversals. The mean of the positive (normal polarity) ChRM directions in the Huelva-1 core is 33.6° ± 17.5°, whereas the mean of the negative (reversed polarity) ChRM directions is -39.4° ± 14.5° (the error given is the standard deviation because the azimuth of the boreholes is unknown and α₉₅ cannot be calculated). For the Montemayor-1 core, these mean values are 30.6° ± 15.8° and -37.9° ± 12.9°, respectively. The mean values are significantly shallower than the expected inclination for a geocentric axial dipole field at the studied site latitude during the late Miocene (over 50°). The azimuth of the boreholes is unknown, therefore the magnetic polarity stratigraphy is based on the ChRM inclination. Eleven polarity intervals are documented (labeled N1–N5 and R1–R6 for normal and reversed polarity intervals, respectively). The Huelva-1 core includes magnetozones R1–R5, whereas the Montemayor-1 core includes magnetozones R1–R6. The polarity record has a largely square-wave shape that attests to its quality. The exceptions are magnetozones N3 and R4 in the Huelva-1 borehole (Figure 3), where reversed and normal polarity directions are recorded down to 2.7 and 4.6 m below their upper boundaries, respectively. A sharp decrease in NRM intensity occurs at around 208 m and 76 m in the Montemayor-1 and Huelva-1 boreholes, respectively (Figures 3 and 4).

Planktic foraminiferal event 2 (appearance of abundant Globorotalia menardii, dextral coiling) is identified in the uppermost part of magnetozone R1 in the Huelva-1 borehole (Figure 3). PF3 (regular appearance of G. miotumida, a marker for the Tortonian/Messinian boundary) and PF4 (dextral coiling of Neogloboquadrina acostaensis) are identified within magnetozone R2 and at the top of magnetozone N3, respectively, in both boreholes (Figures 3 and 4). PF 5 (disappearance of G. miotumida) is only identified in the Huelva-1 borehole, in the uppermost part of magnetozone R4. PF6 (first abundant occurrence of G. margaritae) has been identified in both boreholes near the magnetozone N4/R5 boundary. The appearance of G. puncticulata is found in the lower part of magnetozone R6 in the Montemayor-1 core.

The magnetozone pattern and the position of PF events enable straightforward correlation of the studied cores to the astronomically tuned geomagnetic polarity timescale (ATNTS2004) of Lourens et al. (2004; Figures 3 and 4). PF2, PF3, and PF4 are located within chrons C3Br.2r, C3Br.1r and the lower part of C3An.1r, respectively (Lourens et al., 2004). Thus, magnetozone R1 must correlate with chrons C3Br.2r, R2 with C3Br.1r and R4 with C3An.1r. This correlation implies that the long magnetozone R5 correlates with C3r, which is consistent with the long duration of this chron and with the presence of PF6 near the boundary with chron C3An.1n. The appearance of G. puncticulata is dated at 4.52 Ma near the top of chron C3n.2n (Lourens et al., 2004). Keeping in mind the litoral facies of the Arenas de Huelva Formation and the possible lack of continuity in such facies, this suggests that magnetozone N5 represents an amalgamation of the lower part of chron C3n due to discontinuous sedimentation (Jiménez-Moreno et al., 2013). In this case, magnetozone R6 probably correlates with chron C3n.1r. The magnetobiostratigraphic results presented here indicate that the Huelva-1 borehole is continuous and spans from the latest Tortonian (chron C3Br.2r, ca. 7.4 Ma) to the latest Messinian (uppermost part of chron C3r, ca 5.4 Ma). The Montemayor-1 borehole spans from the latest Tortonian (chron C3Br.2r, ca 7.4 Ma) to the lower Pliocene (Zanclean, chron C3r/C2Ar boundary, ca 4.3–4.2 Ma), which provides a continuous late Miocene record. The only stratigraphic gap appears to be at the base of the Huelva Formation, which marks a sedimentary break that is attributed to a combination of sea level changes (González-Delgado et al., 2004) and tectonic processes (Salvany et al., 2011). Oxygen isotopic data for benthic and planktic foraminifera from the Montemayor-1 borehole enable identification of three distinctive glacial-interglacial stages (TG7, 12 and 22) within chron C3r, which enable refinement of the age model within this long chron (Pérez-Asensio et al., 2012a; Jiménez-Moreno et al., 2013). The age models thereby established for the Huelva-1 and Montemayor-1 boreholes enable calculation of linear sediment accumulation rates (SARs), which have similar trends at both boreholes (Figures 3 and 4). SARs were < 5 cm/kyr until around 6.4 Ma with the exception of a transient increase (up to 8–14 cm/kyr) near the Tortonian/Messinian boundary. After 6.4 Ma, SARs progressively increased and underwent a major rise (exceeding 15 cm/kyr) at around 6 Ma, broadly coincident with the onset of the MSC.

Our rock magnetic study focuses on the clays of the Gibraleón Formation because they represent most of the recovered sequences and, as opposed to the Huelva and Niebla formations, have high NRM intensities and sedimentary facies suitable for hosting biogenic magnetite. The main decay observed in the χ–T heating curves for samples located below 208 m and 76 m in the Montemayor-1 and Huelva-1 boreholes, respectively, which are characterized by high NRMs, is a drop at around 580°C (Figures 6A,B). This drop, which is sometimes preceded by a Hopkinson peak (Figure 6B), indicates the presence of magnetite. Above 580°C, the magnetic susceptibility signal persists and does not disappear completely until about 680°C (Figures 6A,B), which indicates variable contributions from hematite. In many samples, a subtle hump can be observed superimposed on the χ–T curve just below 300°C (Figures 6A,B). This can be interpreted as due to the thermally induced breakdown of minor amounts of maghemite (Liu et al., 2012). Cooling curves typically have higher χ values than heating curves below 580°C, which indicates widespread formation of magnetite during heating (Figures 6A,B). In samples above 208 m and 76 m in the Montemayor-1 and Huelva-1 boreholes, respectively, which are characterized by lower NRMs, thermally induced formation of magnetite above 450°C completely obscures the χ signal (Figure 6C). In these cases, a major increase in χ in the cooling curve below 530°C might suggest that the newly formed magnetite is not stoichiometric.

Hysteresis loops for all of the studied samples saturate below 200 mT, which is consistent with the dominance of magnetite. All samples lie between the SD field and the upper left-hand part of the pseudo-single domain (PSD) region of the Day plot (Day et al., 1977; Figure 7). Samples lie slightly to the right of the theoretical line that represents mixtures of about 60–100% of equidimensional SD magnetite with multi-domain (MD) magnetite (Dunlop, 2002), which suggests an additional contribution from superparamagnetic (SP) material. The mean M_r/M_s value for Huelva-1 and Montemayor-1 samples is 0.37 and 0.43, respectively, which is close to the theoretical value of 0.5 for uniaxial SD magnetite grains and similar to other sediments where biogenic SD magnetite dominates the magnetic mineral assemblage (Kopp et al., 2007; Lippert and Zachos, 2007; Roberts et al., 2011; Larrasoaña et al., 2012; Roberts et al., 2012, 2013). Samples from the upper parts of the two boreholes, where NRM intensities are distinctively lower, are characterized by larger M_r/M_s values (Figure 7), which points to a larger relative contribution of SD magnetite in these sediments.

First-order reversal curve distributions of samples from the two studied boreholes are similar in all cases. They are characterized by closed concentric contours about a central peak located between 20 and 30 mT (Figure 8). Vertical profiles through the peak of the coercivity distributions have a narrow distribution (e.g., <5 mT) around the dominant central peak, whereas horizontal profiles along B_b = 0 mT have a skewed Gaussian shape with variable spreading. These features, which are statistically significant at the 0.05 level (see dark contours calculated following Heslop and Roberts (2012) in Figure 8), define the “central ridge” signature that is typical of intact magnetosome chains (Egli et al., 2010; Roberts et al., 2011, 2012, 2013). Variable coercivity peaks likely indicate mixtures of the so-called “biogenic soft” and “biogenic hard” magnetosome components, whose coercivities (of about 20 and 40 mT, respectively) likely reflect different magnetosome morphologies (Egli et al., 2010; Yamazaki and Ikehara, 2012; Roberts et al., 2013). In some cases (Figures 8A,C), the slight spreading of the outer contours around the peak of the FORC distributions points to some magnetostatic interactions due likely to disruption of a portion of the total number of magnetosome chains (Egli et al., 2010; Kind et al., 2011; Li et al., 2012; Roberts et al., 2012, 2013). The background magnetization observed in the vertical profiles below the narrow central peak suggests an additional contribution of coarser-grained magnetite of probable detrital origin. This detrital component appears to also be responsible for the large spreading observed in the horizontal profile in some samples (Figure 8B). The overall weaker magnetization of samples from the upper parts of boreholes Huelva-1 and Montemayor-1 made high-resolution FORC measurements too noisy to enable reliable identification of magnetic components. Nevertheless, hysteresis ratios of weakly magnetic samples are similar to those of samples with a clear central ridge signature in the FORC diagrams, which suggests that magnetofossils also dominate their magnetic assemblages.

Low-T data indicate the presence of the Verwey transition at around 105 K both in the FC and ZFC curves in all samples (Figures 6D–F). The FC and ZFC cooled curves typically diverge below the Verwey transition, which gives further support for the occurrence of magnetofossils in the studied samples. δ_FC/δ_ZFC values range between ~1 and 1.21, which suggests that the magnetosome surfaces are partially oxidized (Moskowitz et al., 1993). This is consistent with the occurrence of maghemite as evidenced by the χ–T curves (Figures 6A–C).

Scanning electron microscope observations indicate that the studied sediments contain minor amounts of pyrite that occur typically within foraminiferal tests or near other organic remains such as coccolithophore shells (Figure 9). Most pyrite occurs in framboids that are <15 μm in diameter and that consist of small (<1 μm) individual crystals. Framboidal pyrite is often accompanied by euhedral pyrite crystals that are typically <15 μm in size (Figure 9A).

χ–T and low-T experiments, coupled with hysteresis and FORC results, indicate that the magnetic mineral assemblage of the Gibraleón Formation clays is dominated by fossilized intact magnetosome chains that are in some cases partially oxidized. Significant disruption of magnetofossil chains might have been prevented by adhesion of magnetofossil chains onto the surface of clay particles (e.g., Mao et al., 2014). In the absence of TEM images, an alternative origin for fine-grained magnetite linked to diagenetic oxidation of pyrite might be proposed (e.g., Brothers et al., 1996; Rowan and Roberts, 2006). We discard this possibility because: (1) pyrite in the studied clays is never observed in association with iron oxide overgrowths (Figure 9), and (2) this process often occurs during late diagenesis and would result in a prominent paleomagnetic overprint (Brothers et al., 1996; Rowan and Roberts, 2006) rather than in a pristine magnetostratigraphic record (Figures 3 and 4). Nevertheless, TEM observations are needed to corroborate the occurrence of magnetofossils in the studied clays and to assist in identifying different magnetosome morphologies. The Gibraleón Formation clays also contain variable amounts of hematite, which is a common constituent of Saharan dust transported into the NE Atlantic Ocean and has been reported in Quaternary sediments off the SW Iberian margin (Channell et al., 2013b). An eolian origin is therefore most likely for the hematite in the studied clays, although its contribution to the NRM is negligible (Figure 5). The NRM intensity of the clays is linked to the concentration of magnetofossils, as indicated by the similar NRM and M_s variations in both studied boreholes (Figure 10). This indicates that the magnetofossil concentration affects the quality of paleomagnetic data by enhancing the NRM. The mean of both normal and reversed polarity directions for the two studied boreholes is 15° to 20° shallower than expected. Inclination shallowing in fine-grained marine sediments has been typically interpreted as due to rolling of magnetic particles as they are deposited on a substrate for a DRM (e.g., King, 1955) or to sediment compaction that will affect a PDRM carried by detrital magnetic particles (Kent, 1973). Given that the magnetic mineral assemblage in the studied sediments is dominated by intact magnetosome chains, which are typically elongated with a length/width ratio that exceeds five (see Faivre and Schüler, 2008), a likely explanation for the observed paleomagnetic inclination shallowing is depositional flattening of magnetosome chains into the bedding plane. Subsequent sediment compaction might have also contributed to the observed inclination flattening, especially if magnetosome chains are adhered onto clay particles (Mao et al., 2014). More detailed studies are necessary to separate the contribution from these processes, which likely operate in concert in clay-rich lithologies that contain intact magnetofossil chains. Methods that enable discrimination between subfabrics of paramagnetic (e.g., low-T AMS) and ferrimagnetic minerals (e.g., AARM) would be of special relevance. Regardless, magnetofossils from the Gibraleón Formation are interpreted to carry a BDRM, which has recorded all polarity chrons regardless of their short durations, especially near the Tortonian/Messinian boundary (e.g., <50 kyr; Figures 3 and 4). Possible delayed magnetizations have only been found down to 2.7 and 4.6 m below the uppermost boundaries of chrons C3An.1r and 2n, respectively, in the Huelva-1 borehole (Figure 3). Keeping in mind SARs of about 4 and 8 cm/kyr for these chrons, respectively (Figure 10), a delay of 68 and 58 kyr, respectively, can be estimated for the lock in time of magnetizations carried by magnetofossils at these depths. Higher-resolution studies are necessary to confirm and better characterize any BGRM in the studied boreholes.

Benthic foraminiferal assemblages and sedimentary facies indicate that clays of the Gibraleón Formation accumulated in a slope environment, although its lowermost (until about 7.2 Ma) and uppermost (after about 5.85 Ma) parts (Figure 10) accumulated in outer continental shelf environments (Pérez-Asensio, 2012; Pérez-Asensio et al., 2012b). Benthic foraminiferal assemblages point to oligotrophic and mesotrophic conditions during deposition of most of this formation, with the exception of one short-lived interval before 7.2 Ma (Figure 10; Pérez-Asensio, 2012). Oxygen isotopic data indicate that the lower GB was flushed by both the MOW and the AUW until 6.18 Ma (Pérez-Asensio, 2012; Pérez-Asensio et al., 2012a). After 6.18 Ma, rapid reduction and cessation of flow of the MOW suggests closure of the last marine Betic Corridors (Pérez-Asensio et al., 2012a; Figure 10). Later, at about 5.97 Ma, the onset of the MSC led to disconnection of the Mediterranean basin from the Atlantic Ocean (Manzi et al., 2013). Our results indicate that magnetofossils dominate the magnetic mineral assemblage throughout the Gibraleon Formation regardless of changes in depositional setting, nutrient conditions, SARs, and important paleoceanographic events such as cessation of flow of the MOW and onset of the MSC (Figure 10). Mild, suboxic diagenetic conditions are required for preservation of magnetofossils through expanded sections of pelagic sediments that record a syn-depositional remanent magnetization (Roberts et al., 2011, 2013). Clays from the Gibraleón Formation are greenish-bluish in colour, which suggests that iron has been reduced and that sulphate reduction might have occurred. It should be noticed, however, that χ–T and low-T experiments indicate that magnetosomes are partially converted into maghemite by surficial oxidation, and it is unlikely that maghemite would have survived reductive dissolution (Smirnov and Tarduno, 2000). SEM observations of pyrite microtextures indicate that reducing diagenetic conditions in clays from the Gibraleón Formation were mild and closely linked with organic-rich microenvironments (e.g., within foraminiferal tests). We interpret that flushing of the bottom waters in the lower GB by either the MOW and/or the AUW, coupled with mostly oligotrophic and mesotrophic conditions, led to mild diagenetic conditions suitable for flourishing of MTB and preservation of their magnetofossils throughout the Gibraleón Formation. Our results expand the settings in which magnetite magnetofossils can dominate the magnetic properties of sediments to include continental margin sedimentary sequences.

In view of the mild diagenetic conditions that favored preservation of magnetofossils throughout the Gibraleón Formation, and keeping in mind that they largely appear to carry a BDRM that is affected by inclination shallowing, we interpret that MTB lived within the surface mixed layer and, therefore, carry a reliable syn-depositional paleoenvironmental signal. In this case, the main feature needing explanation is the sharp drop observed in magnetofossil abundance at 5.9 Ma and 5.97 Ma in the Huelva-1 and Montemayor-1 boreholes, respectively (Figure 10). Pollen data from the Montemayor-1 borehole indicate that glacial/interglacial variability reported at orbital timescales was not significantly modified after the onset of the MSC at 5.97 Ma (Jiménez-Moreno et al., 2013), which suggests that climate variability is not the underlying cause for the observed drop in magnetofossil abundances. In both boreholes, the decreased magnetofossil content predates the change from meso- to oligotrophic conditions and from slope to shelf environments at 5.85 Ma, and postdates cessation of flow of the MOW by more than 200 kyr (Figure 10). In the Montemayor-1 borehole, however, the sharp shift in magnetofossil abundance coincides strikingly with the onset of the MSC at 5.97 Ma (Manzi et al., 2013; Figure 10B). It should be noted that the age of the drop was established by assuming a linear SAR between the two nearest tie points, namely the chron C3r/C3An boundary and glacial stage TG22 (Figure 10B). For the Huelva-1 borehole, the relevant tie points are the chron C3r/C3An boundary and the base of the Huelva Formation, which is located about 20 m higher in the borehole and, according to results from Montemayor-1, is dated near the Miocene/Pliocene boundary (Figure 4). Keeping in mind this coarser age constraint, we interpret that the sharp drop in magnetofossil abundances is synchronous at the two boreholes and coincident with the MSC onset. The estimated SARs are also constrained by the positions of tie points used to construct the age models. We therefore associate the drop in magnetofossil abundance to a major rise in SAR at the chron C3r/C3An boundary (6.033 Ma), which follows a steady increase that started at 6.5 Ma (Figure 10). This link between magnetofossil abundance and SAR is supported by the overall inverse correlation between M_s and SAR (Figure 11), and therefore suggests that magnetofossil concentrations are controlled by terrigenous dilution. High-frequency changes in NRM intensity also suggest changes in magnetofossil abundance over much shorter timescales, but the extent and significance of such changes need to be assessed with more detailed rock-magnetic data. The marked rise in SAR appears to be linked to the onset of the MSC at 5.97 Ma (Manzi et al., 2013). We tentatively attribute this increased SAR to a change in sedimentation pattern in the GB during the MSC. Thus, closure of the Betic Corridors at that time might have conditioned delivery of large amounts of sediment that resulted from denudation of the Betic Cordillera, which was previously deposited in the Mediterranean Sea, into the Atlantic Ocean. This, in turn, would have accelerated westward progradation of sedimentary systems along the axis of the GB and the arrival of enhanced terrigenous inputs to the locations of the studied boreholes. Regardless of this interpretation, our results indicate that magnetofossils from the Gibraleón Formation carry paleoenvironmental signals that will help to better constrain the paleoceanographic and sedimentary evolution of the GB during the MSC.