# MOW strengthening and contourite development over two analog climate cycles (MIS 12–11 and MIS 2–1) in the Gulf of Cadíz: An impact on North Atlantic climate during deglaciation V and MIS 11?

Paul Moal-Darrigade, Emmanuelle Ducassou, Jacques Giraudeau, André Bahr, Stefanie Kaboth-Bahr, Vincent Hanquiez, Marie-Claire Perello

<figure><img src="image_1.png" /></figure>

## To cite this version:

Paul Moal-Darrigade, Emmanuelle Ducassou, Jacques Giraudeau, André Bahr, Stefanie Kaboth-Bahr, et al.. MOW strengthening and contourite development over two analog climate cycles (MIS 12–11 and MIS 2–1) in the Gulf of Cadíz: An impact on North Atlantic climate during deglaciation V and MIS 11?. Global and Planetary Change, 2022, 208, pp.103721. 10.1016/j.gloplacha.2021.103721. hal-04285816

HAL Id: hal-04285816

https://hal.science/hal-04285816v1

Submitted on 25 Jul 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_2.png" /></figure>
1 MOW strengthening and contourite development over two analog

2 climate cycles (MIS 12–11 and MIS 2–1) in the Gulf of Cadíz: an

3 impact on North Atlantic climate during deglaciation V and MIS

4 11 ?

Paul Moal-Darrigadeª, Emmanuelle Ducassouª, Jacques Giraudeauª, André Bahrᵇ, Stefanie Kaboth-Bahrᶜ, Vincent Hanquiezª, Marie-Claire Perelloª

7

8 Abstract

9

10 High-resolution records from International Ocean Discovery Program (IODP) sites

11 U1386 and U1387 drilled during IODP Expedition 339 into the Faro Drift, made it possible to

12 assess the impact of intensifications of the upper core (MOWU) of the Mediterranean Outflow

13 Water (MOW) and of changes in sediment supply on the sedimentation in the northern Gulf

14 of Cadíz since the Middle Pleistocene. This work focuses on the comparison of records

15 covering Marine Isotope Stage (MIS) 2–1 and MIS 12–11, in order to investigate the

16 behaviour and circulation regime of the MOWU over two climatic cycles of similar

17 astronomical configurations and their associated deglaciation. The analysis of facies

<footnote>ª Université de Bordeaux, UMR CNRS 5805 EPOC, Allée Geoffroy St Hilaire, 33615 Pessac, France</footnote>

<footnote>ᵇ Institute of Earth Sciences, Heidelberg University, Im Neuenheimer Feld 234, 69120 Heidelberg, Germany</footnote>

<footnote n="5">5 c Institute of Geosciences, University of Potsdam, Karl-Liebknecht-Straße 24-26, 14476 Potsdam-Golm,</footnote>

<footnote n="6">6 Germany</footnote>
established on the basis of grain-size, XRF core-scanning, and carbonate content revealed
contourite beds formed by the MOWU during MIS 11 and MIS 1 and deglaciations
(deglaciation V and I). Contourite sequences show that MOWU velocity at the seabed was
higher during MIS 2–1 than during MIS 12–11, and that sediment supply was different
between these two climatic cycles. While overall low during MIS 12–11, MOWU intensity
increased during deglaciation V and MIS 11 and preceded large ice rafted events and cooling
in the North Atlantic Ocean. As a major element of the MOW, MOWU strengthening during
deglaciation V likely contributed to higher heat and moisture transport towards the high
latitudes inducing a slight increase of calving and size of boreal ice sheets. The MOW-derived
injection of heat and salt in the North Atlantic Ocean during deglaciation V might have
contributed, through reactivation of the upper AMOC, to the switch of the Atlantic
thermohaline circulation from a glacial to an interglacial mode.

Keywords 30

31 • Gulf of Cadíz

32 • IODP Sites U1386-U1387

33 • Mediterranean Outflow

34 • MIS 12–11 vs. MIS 2–1

35 • Contourites

36 • Grain-size

37

1.Introduction 38

39
Contourite drifts are defined as sediment bodies characterized by a major accumulation of sediment deposited or reworked by a bottom current (Heezen et al., 1966; Stow et al., 2002; Rebesco et al., 2008, 2014). The high sedimentation rates of these drifts make them precious archives for high resolution paleoceanographical studies (e.g., Brunner, 1986; Robinson and McCave, 1994; Harris et al., 2001).

Located off the southwestern sector of the Iberian margin, the Gulf of Cadíz is one of the most studied Contourite Depositional System (CDS) in the world ocean (e.g., Faugères et al., 1984; Hernández-Molina et al., 2006; Mulder et al., 2006; Hanquiez et al., 2007; Marchès et al., 2007; Llave et al., 2007; Brackenridge et al., 2013; Stow et al., 2013; Alonso et al., 2016; Mestdagh et al., 2019). Its morphology is essentially shaped by the activity of the Mediterranean Outflow Water (MOW). Strongly influenced by the African monsoon system, MOW contributed to propagate (sub)tropical climate signals into the high latitudes during the Late Pleistocene (Bahr et al., 2015; Voelker et al., 2015), impacting the Atlantic Meridional Overtuning Circulation (AMOC) (Rogerson et al., 2006; Voelker et al., 2006; Bahr et al., 2015).

The Contourite Depositional System of the Gulf of Cadíz, has been extensively studied for high resolution reconstruction of MOW dynamics over the last 50 kyrs (Sierro et al., 1999; Rogerson et al., 2005; Llave et al., 2006; Voelker et al., 2006; Toucanne et al., 2007). Many studies support an enhancement of the lower branches of the MOW (MOW_L) during cold periods of the last glacial-interglacial cycle, synchronously with abrupt climate changes in the North Atlantic (e.g., Sierro et al., 1999; Schönfeld and Zahn, 2000; Voelker et al., 2006). Some authors also observed a strong upper branch of the MOW (MOW_U) during cooling events in the Northern Hemisphere (Llave et al., 2006; Toucanne et al., 2007; Bahr et al., 2015; Sierro et al., 2020). However, it has been sown that MOW_U only shows strengthening during warm intervals (Llave et al., 2006; Kaboth et al., 2016; Sierro et al.,
2020). Sedimentological analyses, suggest that the composition of sediments in the Gulf of Cadíz is only partly related to changes in bottom current velocity, as changes in sediment supply by rivers can also play an important role in sediment distribution (Mulder et al., 2013). The recent recovery of long sediment sequences from the Faro Drift during IODP Expedition 339 (Hernández-Molina et al., 2014) offers a unique opportunity to extend the chronological framework of MOW reconstruction to older climatic cycles of the Middle Pleistocene, and to test the reliability of sedimentological parameters as proxies of current strength in the Gulf of Cadiz.

The present high resolution sedimentological and micropaleontological study is based on two Faro Drift sediment records recovered at IODP sites U1386 and U1387, both covering, the Marine Isotope Stages (MIS) 2-1 (~ last 29 kyrs) and MIS 12-11 (~478-374 kyrs BP) intervals. These two sites are located 4 km apart and thus provide insights into the potential spatial variability of sedimentation in contourite drifts. MIS 2-1 and MIS 12-11 have been chosen for their similarities in Earth orbital configuration (Loutre and Berger, 2003). Despite differences in deglaciation trajectories and slight discrepancies in orbital configuration (i.e. in the phasing of precession and obliquity) between MIS 11 and MIS 1 (e.g., Loutre and Berger, 2000; Rohling et al., 2010), their common weak eccentricity forcing and resulting similar insolation signals led to consider these two interglacials as close analogues (Loutre and Berger, 2000, 2003). This analogy also refers to comparable MIS 11 vs. pre-industrial MIS 1 greenhouse gas concentrations (CO₂ and CH₄), global atmospheric and ocean temperatures, global mean sea-level, and local sea-surface conditions (e.g., Petit et al., 1999; Droxler and Farrell, 2000; McManus et al., 2003; de Abreu et al., 2005; Raynaud et al., 2005; Siegenthaler et al., 2005; Spahni et al., 2005; Jouzel et al., 2007; Bowen, 2010). Hence MIS 11 has often been considered as the best analogue to the Holocene among the recent interglacials (e.g., Droxler and Farrell, 2000; Loutre and Berger, 2003; de Abreu et al., 2005; Palumbo et al.,
2019). Some studies based on sediments of the Portuguese margin, close to our study area, evidence similar local sea-surface conditions and dynamics during these interglacials (de Abreu et al., 2005; Rodrigues et al., 2011; Palumbo et al., 2019). MIS 12, which is considered as one of the most extreme glacial stage of the last 500 kyrs, shows very similar millennial time-scale variations than those characterizing the last glacial period -MIS 2- (McManus et al., 1999). Moreover, global atmospheric temperatures and Western Iberian sea-surface conditions were comparable during both glacials (e.g., Jouzel et al., 2007; Rodrigues et al., 2011; Palumbo et al., 2019).

The MIS 2–1 interval is the best studied glacial-interglacial cycle albeit MIS 1 does not represent an entire interglacial. Comparing the MIS 12–11 cycle to the MIS 2–1 interval serves as a test for previously established models of MOW dynamics (e.g Llave et al., 2006). We make use of grain-size characteristics and the geochemical composition of sediments to define distinct facies with the objective of understanding the impact of bottom current velocity and sediment supply changes on sedimentation patterns during these two comparable periods. Then, using our record of MOWU activity during MIS 11, we evaluate the possible connection between MOW and the North Atlantic Climate during an analogue of the current interglacial (Loutre and Berger, 2000, 2003).

107

108

## 2. Regional setting

109

110

111

112

The present-day circulation in the Gulf of Cadíz is driven by exchanges through the Strait of Gibraltar where relatively low density (less saline) North Atlantic Water penetrates into the Mediterranean Sea at surface, while denser Mediterranean Water flows into the
113

Atlantic Ocean at depth (e.g., Bryden and Stommel, 1984; Bryden et al., 1994; Baringer and Price, 1999; Naranjo et al., 2015). The MOW which circulates northward following the continental slope (Ambar and Howe, 1979; Hernández-Molina et al., 2006, 2014; Mulder et al., 2006) is the principal water mass at intermediate depths in the Gulf of Cadíz. This saline Mediterranean water is located above the North Atlantic Deep Water (Baringer and Price, 1999; Hernández-Molina et al., 2014) and is overlaid by the sub-polar variety of North Atlantic Central Water (Voelker et al., 2015b).

114

115

116

117

118

119

120

121

122

123

124

125

126

127

128

129

<figure><img src="image_3.png" /><figcaption>Fig. 1: A) Position of ODP Site 1061. B) General currents and water masses influencing the Gulf of Cadíz and location of deep-sea cores MD99-2343 and MD01-2472. C) IODP Sites U1386 and U1387 locations and major water masses of the Gulf of Cadíz (Demirov and Pinardi, 2007; Carracedo et al., 2016; Hernández-Molina et al., 2016; Schroeder et al., 2017)</figcaption></figure>

The MOW, which flows at a rate of about 1.78 Sv through the Strait of Gibraltar (Bryden and Stommel, 1984; Bryden et al., 1994), is principally composed of Levantine Intermediate Water (LIW) produced in the Eastern Mediterranean Sea (Millot et al., 2006; Millot, 2009, 2014) and of variable proportions of Western Mediterranean Deep Water
130

formed in the Gulf of Lion and in the Alboran and Tyrrhenian Seas (Millot, 2014). LIW is the
dominant water mass composing MOW, inferring that deep and intermediate water formation
131
in the Eastern Mediterranean Sea is a key driver of MOW production (Rogerson et al., 2012;
132
Toucanne et al., 2012; Millot, 2014).

131

132

133

134

Upon leaving the Strait of Gibraltar and entering the Gulf of Cadíz, MOW splits into
135
two main flows under the influence of the complex morphology of the continental slope: the
136
Mediterranean Upper Water (MOWU) and the Mediterranean Lower Water (MOWL). MOWL
137
flows between 800 and 1,400 m depth alongside the middle slope, and is subdivided into three
138
minor branches due to slope morphology (e.g. Madelain, 1970; Ambar and Howe, 1979;
139
Borenäs et al., 2002). Following the base of the upper slope at depths of 400–700 m (Ambar,
140
1983; Ambar et al., 2002; Hernández-Molina et al., 2011, 2014), the MOWU flows northward
141
into the Gulf of Cadíz. MOW velocity gradually decreases northward from 300 cm.s⁻¹ in the
142
Strait of Gibraltar to 10–30 cm.s⁻¹ at the latitude of Cape Saint Vincent (Ambar and Howe,
143
1979). Bottom current velocity is however highly variable locally due to seafloor
144
irregularities (Hernández-Molina et al., 2006; García et al., 2009; Stow et al., 2009; Serra et
145
al., 2010)

135

136

137

138

139

140

141

142

143

144

145

146

The Contourite Depositional System of the Gulf of Cadíz results from the complex
147
interaction between bottom water currents of the MOW and the middle slope of the southwest
148
Iberian margin (Madelain, 1970; Faugères et al., 1984; Nelson and Maldonado, 1999; Stow et
149
al., 2002; Mulder et al., 2006; Hernández-Molina et al., 2006; Hanquiez et al., 2007; Marchès
150
et al., 2007). The generated drift deposits, such as the Faro Drift studied in this paper, are
151
composed of different proportions of muddy, silty and sandy sediment of mixed terrigenous
152
and biogenic composition (e.g., Gonthier et al., 1984; Stow et al., 1986).

147

148

149

150

151

152
A series of papers based on the study of CDS sediments demonstrate the relation between the behaviour of the MOW and different climatic events during the late Quaternary (e.g. Nelson and Maldonado, 1999; Schönfeld and Zahn, 2000; Rogerson et al., 2005; Llave et al., 2006; Voelker et al., 2006; Toucanne et al., 2007; Bahr et al., 2014, 2015; Kaboth et al., 2017). It appears that during the Pleistocene, the African monsoon had a strong influence on MOW variability on orbital-scales (Bahr et al., 2015; Voelker et al., 2015; Sierro et al., 2020). On glacial-interglacial time scales, the multiple branches of the MOW occupy a different spatial positions and a fluctuate in their flow strength during warm vs. cold climatic periods due to varying densities of the outflowing water masses (Schönfeld and Zahn, 2000; Rogerson et al., 2005; Hernández-Molina et al., 2006; Llave et al., 2006; Voelker et al., 2006; Rogerson et al., 2012). Sediments in the Gulf of Cadíz have recorded further modifications of the character of the MOW linked to the influence of North Atlantic climate oscillations, with a strengthened MOW during Heinrich Stadials (HS) and Dansgaard-Oeschger (DO) stadials (Llave et al., 2006; Voelker et al., 2006; Toucanne et al., 2007). These findings are supported by models which show increased convection in the North Atlantic when the MOW is the strongest (Bigg and Wadley, 2001), even if the exact impact of the MOW on AMOC remains unclear (Swingedouw et al., 2019).

## 3. Material & Methods

### 3.1. IODP Sites U1386 and U1387

This study is based on two sites drilled during IODP Expedition 339 (Stow et al., 2013) from November to January 2011/2012. IODP sites U1386 and U1387 lie off the
southern Iberian margin, at 36°49.685′N, 7°45.321′W and 561 m depth, and 36°48.321′N, 7°43.1321′W and 560 m water depth, respectively (Fig. 1). Both IODP sites are presently bathed by the upper branch of the MOW but Site U1386 is positioned closer to the moat generated by the MOWU (Hernández-Molina et al., 2013). Sediment samples were selected following the shipboard splice of IODP Sites U1386 and U1387 which joins core sections from Holes A and B (Expedition 339 Scientists, 2012). The high-resolution XRF scanning and grain-size analyses of the sediment from the deepest interval of IODP Site U1386 suggests that the shipboard splice in this depth interval based on physical property correlations (Expedition 339 Scientists, 2013) required modifications. In more detail, we find that the previous core-to-core transition (A15-B14) divided a contourite layer in the middle, that we corrected according to XRF scanning records and grain-size measurements (Fig. 2). The corrected meters composite depth (c-mcd) scale used in this paper presents a correction of + 0.17 m between the core U1386A-15H and U1386B-14H. This paper focuses on the intervals which correspond to 0–10 meters composite depth (mcd) and 126–162 (c-mcd) for site U1386 and 0–9 mcd and 95–125 mcd for site U1387.
<figure><img src="image_4.png" /></figure>

Fig. 2: Correction of the shipboard meter composite depth (mcd) scale of core B_14H (black) and core A_15H (grey) using several parameters including a) Magnetic susceptibility (MS), b) median grain size (D₅₀), and c) XRF derived Fe/Ca ratio. + 0.17 m corresponds to the length of sediment column added to the splice.

## 3.2. Radiometric analyses of IODP Site U1386

Seven radiocarbon measurements were performed on mix, visually clean, handpicked planktic foraminiferal tests from the > 150 µm fraction, according to Ducassou et al. (2018). Ages were determined at the Laboratoire de Mesure du Carbone 14 Saclay (Paris). The calibration of radiocarbon age to calendar years was performed using the Calib Rev. 7.0.2 program/ Marine20 data set (Reimer et al., 2009). Here, we chose to present the median probability of the ages (Telford et al., 2004).
206 3.3. Foraminiferal analyses

207

208 The biostratigraphy of IODP Site U1386 was performed at a resolution close to 25 cm
209 for both intervals. The resolution of faunal analyses for IODP Site U1387 was higher (20 cm)
210 in the upper interval than in the lower one (41 cm). The samples were washed and sieved in
211 order to extract the > 63µm fraction. The residues were dried and weighed and the > 150 µm
212 fraction was separated for foraminiferal counts.

213 Planktic foraminiferal identification was performed in aliquots of the > 150 µm
214 fraction with total numbers of specimens exceeding 250, and according to Spezzaferri et al.
215 (2015) and Schiebel and Hemleben (2017). In this study, special attention was directed
216 towards the species *Negloboquadrina pachyderma*, *Globorotalia truncatulinoides*,
217 *Globigerinoides ruber rosea*, *Globogerinoides conglobatus*, *Globorotalia crassaformis* and
218 *Trilobatus sacculifer*. Species abundances were expressed in weight % according to total
219 counts.

220 The abundance of the polar species *N. pachyderma* was used for tracing the presence
221 of subpolar to polar surface waters in the southwestern Iberian margins during cold events of
222 Heinrich Stadials (e.g., Eynaud et al., 2009; Voelker et al., 2009; Ducassou et al., 2018).

223 *G. ruber rosea*, *G. conglobatus*, *G. crassaformis* and *T. sacculifer* are tropical and sub-
224 tropical species and their abundance in sediments of the Gulf of Cadiz increased during the
225 recent warm periods of the Holocene and the Bølling-Allerød (e.g., Sierro et al., 1999;
226 Ducassou et al., 2018).

227 *G. truncatulinoides* is a subtropical deep dwelling species and is a good indicator of
228 the thermocline location (e.g., Be, 1960). Coiling direction of *G. truncatulinoides* are given as
229 percentages of the sinistral coiling ratio, using the formula: GTS =GTS * 100 / (GTS + GTD),
where GTS is the number of specimens of *G. truncatulinoides* sinistral and GTD the number of specimens of *G. truncatulinoides* dextral (Ducassou et al., 2018). This coiling ratio was calculated only when the number of specimens represented at least 0.5 % of the total assemblage.

234

## 3.4. X-ray fluorescence analysis

236

High-resolution X-ray fluorescence (XRF) scanning was conducted along the split core of U1386, from 125 to 165 mcd, using an Avaatech XRF corescanner III at Center for Marine Environmental Sciences (MARUM, University of Bremen). The XRF III scanner is equipped with an Oxford Instrument 100W Neptune X-ray tube with Rhodium (Rh) target. Measurements were conducted every 3 cm over a 1.2 cm² area with a down-core slit size of 12 mm and a sampling time of 20 s. Two separate runs were performed using generator settings of 10 kV and 0.4 mA, and 30 kV and 1.0 mA. The split core surface was covered with a 4 micron-thin SPEXCerti Prep Ultralenel foil to avoid sediment desiccation and contamination between XRF measurements. The data were acquired by a Camberra W-PIPS Silicon Drift Detector (SDD; Model SXD 15C-150-500) with 150 eV X-ray resolution and the Canberra Digital Spectrum Analyzer DAS 1000. Raw data were processed by the analysis of X-Ray spectra by the Iterative Least square software (WIN AXIL) package from Canberra Eurisys. The 0-10 mcd interval of IODP Site U1386 was investigated with an AVAATECH XRF Scanner of the Royal Netherlands Institute for the Sea Research (NIOZ) (details given in Kaboth et al. (2016)). XRF measurements of U1387 material were performed at the MARUM-University of Bremen using an Avaatech XRF corescanner II (detail given in Bahr et al. (2014)). XRF core-scanning is a time efficient method for semi-quantitative elemental analysis of sediment. Individual elemental intensities (expressed in counts) can be influenced
by changing lithology and by the formation of a water film below the covering foil (Tjallingii et al., 2007). These biases can be corrected by using elemental ratios (Bahr et al., 2014).

In the current study, we are presenting the ln(Zr/Al) and ln(Fe/Ca) records. The ln (Zr/Al) ratio is a useful semi-quantitative indicator of bottom current velocity variations (Bahr et al., 2014). Furthermore, ln(Fe/Ca) reflects the contribution of biogenic carbonate-rich particles versus detrital components in sediment supply.

261

## 3.5. Sedimentological analyses

263

Grain-size analyses were performed on the bulk sediment, at 5 cm intervals in coarser sequences and at 100 cm (U1386) and 150 cm (U1387) resolutions in homogeneous muddy sequences, using a Malvern MASTERSIZER S (University of Bordeaux, UMR 5805 EPOC).
In the topmost interval of IODP Site U1386 (0–10 mcd), grain-size analyses were also performed on decarbonated samples. Measurements of median grain size ($D_{50}$), percentage of sand and sorted silt mean size (SS; McCave et al. (1995)) were used here as indexes of bottom flow speed according to McCave and Hall (2006). To make up for the lack of decarbonated grain-size measurements on the deepest intervals carbonate content was measured using a Bernard calcimeter between 129 and 162 c-mcd of IODP Site U1386. We collected 12 samples of 50–130 mg on the deepest interval of IODP Site U1386, at different strategic locations chosen according to the ln(Fe/Ca) ratio. In the > 150 µm fraction, lithic grain, including quartz grains, were counted and expressed as number per gram of dry weight of bulk sediment.

277

# 4. Chronostratigraphy of IODP Sites U1386 and U1387
279

280

4.1.Stratigraphy of IODP Site U1386

281

282

283

284

285

286

287

288

289

The initial chronology for both MIS 12–11 and MIS 2–1 intervals at IODP Site U1386 is based on the visual alignment of the benthic stable oxygen isotope record from IODP Site U1386 to the global benthic δ¹⁸O record of LR04 (Kaboth et al., 2017). For the interval MIS 2–1 the initial chronology was refined by six new AMS ¹⁴C dates (Table 1), validating previous age models proposed by Bahr et al. (2015) and Kaboth et al. (2017) for the last 20 kyrs BP and slightly improving the chronology between 30 and 20 kyrs BP (Fig. 3). Based on the biostratigraphic framework and benthic δ¹⁸O record, one of the seven AMS ¹⁴C dates was not considered in the construction of the age model (Table 1 and Fig. 3).

290

<caption>Table 1: New radiocarbon ages of IODP Site U1386 (this study). Bulk: mix of planktic foraminifera.</caption>

291

<table><thead><tr><th>Site</th><th>Depth (mcd)</th><th>Lab code</th><th>Species</th><th>Conventional AMS ¹⁴C age (¹⁴C yr BP)</th><th>Standard error</th><th>95.4% (2 σ) cal yr BP age ranges</th><th>Cal yr BP age median probability</th><th>Remarks</th></tr></thead><tbody><tr><td rowspan="6">U1386</td><td>0.18</td><td>SacA 35774</td><td>Bulk</td><td>2155</td><td>± 30</td><td>1,410–1,707</td><td>1,5</td><td></td></tr><tr><td rowspan="2">1.08</td><td rowspan="2">SacA 35775</td><td rowspan="2">Bulk</td><td rowspan="2">6735</td><td rowspan="2">± 40</td><td rowspan="2">6,870–7,223</td><td>68</td><td></td></tr><tr><td>7,0</td><td></td></tr><tr><td rowspan="2">1.68</td><td rowspan="2">SacA 35778</td><td rowspan="2">Bulk</td><td rowspan="2">8570</td><td rowspan="2">± 40</td><td rowspan="2">8,809–9,217</td><td>44</td><td></td></tr><tr><td>9,0</td><td></td></tr><tr><td rowspan="2">2.38</td><td rowspan="2"></td><td rowspan="2"></td><td rowspan="2"></td><td rowspan="2"></td><td rowspan="2"></td><td>15</td><td></td></tr><tr><td>10,</td><td></td></tr><tr><td rowspan="2">4.438</td><td rowspan="2"></td><td rowspan="2">SacA 35779</td><td rowspan="2">Bulk</td><td rowspan="2">9805</td><td rowspan="2">± 50</td><td rowspan="2">10,350–10,801</td><td>58</td><td></td></tr><tr><td>7</td><td></td></tr><tr><td rowspan="2">6.438</td><td rowspan="2"></td><td rowspan="2">SacA 35776</td><td rowspan="2">Bulk</td><td rowspan="2">20170</td><td rowspan="2">± 100</td><td rowspan="2">23,012–23,674</td><td>23,</td><td>Assumed to be too old based on benthic δ¹⁸O and biostratigraphic framework</td></tr><tr><td>32</td><td></td></tr><tr><td></td><td></td><td>SacA 35777</td><td>Bulk</td><td>23060</td><td>± 140</td><td>26,027–26,864</td><td>26,</td><td></td></tr></tbody></table>
<table><tr><td>13.61</td><td></td><td></td><td></td><td></td><td>56</td></tr><tr><td>SacA 35780</td><td>Bulk</td><td>27230</td><td>± 180</td><td>30,156–30,943</td><td>9</td></tr></table>

292

<figure><img src="image_5.png" /></figure>

Fig.

Fig. 3: A) age model of IODP Site U1386 over the past 40 kyrs cal. B) Age model of IODP Site U1386 between 485 kyrs BP and 365 kyrs BP. Black circles indicates correlation points to northern Atlantic records (see Supplementary Data Fig. 1), black squares represent <sup>14</sup>C-based control points, correlation points to LR04 are marked by light green diamond (Kaboth et al., 2017); dark green circles show correlation points (XRF-Br) to U1389. Tie points of IODP Site U1386 used for this study are indicated in Table 2.

293

294

295

296

297

298

299

300

301

302

303

304

305

The resulting sedimentation rates at the beginning of MIS 2 are very high at IODP Site U1386 (173 cm kyrs⁻¹). The remaining interval of MIS 2 as well as the following deglaciation is characterized by a drop of sedimentation rate to 24.5 cm kyrs⁻¹ (Table 2). During the lower Holocene, rates briefly increase (up to 45 cm kyrs⁻¹) before declining progressively to reach 16 cm kyrs⁻¹ during the upper Holocene (Table 2). This comparatively low sedimentation rate possibly reflects a low supply of sediment during the end of Holocene. According to sampling intervals, these various sedimentation rates result in a mean resolution
306

of 0.46 kyrs between each planktic foraminiferal sample used to infer bioevents of IODP Site U1386.

307

308

<caption>Table 2: Tie points of IODP Site U1386 are based on ¹⁴C dates, alignment of isotopic events of δ¹⁸O benthic record at IODP Site U1386 (during MIS 11b) with δ¹⁸O benthic record from North Atlantic (see Supplementary data Fig. 1) and on a correlation between the δ¹⁸O benthic record (Kaboth et al., 2017) of IODP Site U1386 and the LR04 benthic stack of Lisiecki and Raymo (2005). a Tie point used to correlate IODP Site U1386 record to LR04 benthic stack of Lisiecki and Raymo (2005) (Kaboth et al., 2017). In black: correlation points of this study. In grey: tie point ruled out for this study.</caption>

309

310

311

312

313

<table><thead><tr><th>IODP Site<br>U1386 depth<br>(mcd and c-<br>mcd)</th><th>Tie point age<br>(Kyrs)</th><th>Sed. Rate (cm kyrs<sup>-1</sup>)</th></tr></thead><tbody><tr><td>0.18</td><td>1.568</td><td></td></tr><tr><td>1.08</td><td>7.044</td><td>16</td></tr><tr><td>1.68</td><td>9.015</td><td>30</td></tr><tr><td>2.38</td><td>10.587</td><td>45</td></tr><tr><td>4<sup>a</sup></td><td>18</td><td>22</td></tr><tr><td>6.438</td><td>26.433</td><td>29</td></tr><tr><td>13.613</td><td>30.569</td><td>173</td></tr><tr><td>117.4<sup>a</sup></td><td>341</td><td></td></tr><tr><td>129.6</td><td>383</td><td>29</td></tr><tr><td>133</td><td>388</td><td>68</td></tr><tr><td>136.1</td><td>398</td><td>31</td></tr><tr><td>136<sup>a</sup></td><td>401</td><td></td></tr><tr><td>138.5<sup>a</sup></td><td>405</td><td>34</td></tr><tr><td>144.17<sup>a</sup></td><td>431</td><td>27</td></tr><tr><td>172.17<sup>a</sup></td><td>508</td><td>33</td></tr></tbody></table>

314

315

We also slightly adapted the chronology for the MIS 12—11 time interval by modifying the applied mcd depth scale. During the MIS 12–11 cycle, mean sedimentation rates averaged 34 cm.kyrs⁻¹ (Table 2) and the sampling interval allows for a mean time resolution of ~0.9 kyrs between each planktic foraminiferal sample.

316

317

318
319 4.2SITE-TO-SITE CORRELATION (U1386-U1387)

320

321 4.2.1. Planktic foraminifer distribution: a correlation tool

322

323 A sediment transport and sorting independent proxy based on planktic foraminifer distribution was used as a tuning target for the site-to-site correlation. Hence differences in 324 MOW flow at both sites can be more independently assessed than by using other tuning such 325 as Zr/Al signals. Transferring the age model from our reference IODP Site U1386 into U1387 326 (Fig. 4 and Table 3), is based on the following arguments. First, a recent study of Ducassou et 327 al. (2018) showed that during the Late Pleistocene, bioevents based on planktic foraminifera 328 from Gulf of Cadiz present similar ages regardless of sedimentary environments (contourite 329 vs. hemipelagite). Second, some of the species are clearly related to rapid changes in surface 330 water conditions in the Gulf of Cadiz which provide additional informations on the 331 sedimentation conditions. 332

333

334

335 **Table 3:** Stratigraphic tie points in IODP Site U1387 based on correlations of planktic foraminiferal abundances (N. 336 pachyderma, G. truncatulinoides and G. crassaformis s) with bioevents identified in IODP Site U1386, and XRF-Br 337 correlations to U1389 (Bahr et al., 2015).<sup>b</sup> Tie point used to correlate the XRF-Br of IODP Site U1386 to IODP Site 338 U1389 (Bahr et al., 2015). Sedimentation rate in italic has to be considered carefully because of 339 uncertainty about the age model.

<table><thead><tr><th>IODP Site<br>U1387 depth<br>(mcd)</th><th>Tie point ages<br>(kyrs)</th><th>Sedimentation<br>rate (cm.kyrs<sup>-1</sup>)</th></tr></thead><tbody><tr><td>0</td><td>1.568</td><td></td></tr><tr><td>1.01</td><td>7.442</td><td>17</td></tr></tbody></table>
<table><tr><td>1.9</td><td>10.656</td><td>28</td></tr><tr><td>2.19</td><td>12</td><td>22</td></tr><tr><td>2.5</td><td>16.49</td><td>7</td></tr><tr><td>3.55</td><td>19.4</td><td>36</td></tr><tr><td>7.25</td><td>27.9</td><td>44</td></tr><tr><td>9.91<sup>b</sup></td><td>30.09</td><td>121</td></tr><tr><td>95.7</td><td>374</td><td></td></tr><tr><td>97.8</td><td>383</td><td>23</td></tr><tr><td>102.3</td><td>400</td><td>26</td></tr><tr><td>104.6</td><td>405.5</td><td>42</td></tr><tr><td>105.8</td><td>415</td><td>13</td></tr><tr><td>107.7</td><td>428.5</td><td>14</td></tr><tr><td>109</td><td>430.5</td><td>65</td></tr><tr><td>113.2</td><td>445</td><td>29</td></tr><tr><td>120</td><td>468.5</td><td>29</td></tr><tr><td>121.2</td><td>472</td><td>34</td></tr></table>

340

<figure><img src="image_6.png" /></figure>

<figure><img src="image_7.png" /><figcaption>Fig. 4: A) age model of IODP Site U1387 over the past 40 kyrs cal. B) Age model of IODP Site U1387 between 485 kyrs BP and 365 kyrs BP; Dark green circles show correlation points (XRF-Br) to U1389 (Bahr et al., 2015); brown triangles represent planktic foraminiferal correlation points between U1387 and U1386 and red line represents age models used in this study. Tie points in IODP Site U1387 are given in Table 3.</figcaption></figure>

341

342

343

344

345

346
347

### 4.2.2. From planktic foraminifer assemblages to bioevent identification

348

349

350

351

352

353

354

355

By combining this chronology with the planktic foraminiferal assemblages, we identified 4 and 6 bioevents during the MIS 2–1 and MIS 12–11 periods, respectively. In order to facilitate the description of bioevents and the following discussion, and based on the recent nomenclature recommendation (Railsback et al., 2015), we divide MIS 12 and MIS 11 into three substages (12c, 12b 12a and 11c, 11b and 11a). The MIS 11c interval, which extends from 425 kyrs cal. BP to 398 kyrs cal. BP, corresponds to the interglacial period of the MIS 11 (Tzedakis et al., 2012b).

356

357

### 4.2.2.1. Holocene and MIS 11 foraminiferal assemblages

358

359

360

361

362

G. ruber rosea, G. conglobatus, G. crassaformis sinistral (s) and T. sacculifer were typically present during interglacial periods with abundance ranges of 0–5%, 0–3%, 0–2% and 1–10%, respectively. Our results show an increase of these warm species during the Holocene (Fig. 5) as previously described in the study area (Ducassou et al., 2018).

363

364

365

366

367

368

369

MIS 11 was overall characterized by high contributions (4–16%) of T. sacculifer with peak values reached during substage MIS 11c. The relative abundances of the accessory warm water species G. ruber rosea and G. conglobatus show maximum values (1–2%) during this substage. Rare occurrences (0.3–1.5%) of G. crassaformis (s) during MIS 11c and MIS 11a at both IODP sites U1386 and U1387 (Fig. 6) are used to improve the correlation between both sites when tuning based on the abundance of G. truncatulinoides (see below, section « 4.2.2.3. Coiling direction of *Globorotalia truncatulinoides*») is questionable.

370
371 4.2.2.2.Glacial and cold assemblages: YD, MIS 2 and MIS 12

372

Over the last 30 kyrs, at least three intervals of increased occurrences of the polar species *N. pachyderma* were observed, and were used to correlate IODP Site U1387 to IODP Site U1386. The youngest interval, with mean values of 3%, encompasses the Younger Dryas (e.g., Rogerson et al., 2005; Voelker et al., 2006; Ducassou et al., 2018). In the Gulf of Cadíz, this bioevent was dated between 11,070 and 13,800 cal yr BP (Ducassou et al., 2018). The previous interval of peak occurrence of *N. pachyderma* (11–12%; Fig. 5) is linked to Heinrich Stadial 1 (18 to 15.6 cal kyr BP; Sanchez Goñi and Harrison (2010)) as identified by several authors (e.g., Turon et al., 2003; Ducassou et al., 2018). In the study area, this bioevent extends from 17,950 to 15,400 cal yr BP (Ducassou et al., 2018). The third period of high occurrence of *N. pachyderma* (ca. 3%) is older than 28,400 yr BP. Because of the different characteristics of this third bioevent compared to the two more recent ones, we propose that this final bioevent is related to Heinrich Stadial 3, as previously observed in other studies(e.g., Llave et al., 2006; Ducassou et al., 2018). Alike Voelker et al. (2009) who used *N. pachyderma* abundances in a core of the Gulf of Cadiz (MD99-2399) to reconstruct the hydrography of Iberian margins, we did not find any evidence of Heinrich Stadial 2 in planktic foraminiferal records of IODP sites U1386 and U1387. Abundance of *N. pachyderma* at sites U1386 and U1387 varied between 0 and 16% during MIS 12–11. Peak values of 8 and 16%, respectively, occurred during deglaciationV (Fig. 6) and were used to correlate IODP sites U1387 and U1386. Another correlating bioevent was identified during MIS 12c in the form of a short interval of high contribution of *G. crassaformis* sinistral, with peak values reaching 3% of the total foraminiferal assemblages at both sites.

394
395 4.2.2.3.Coiling direction of Globorotalia truncatulinoides

396

397 Changes in coiling direction of *G. truncatulinoides* are standard bioevents of the
398 Pleistocene. They were initially used to correlate Late Pleistocene cores in the equatorial
399 Atlantic ocean (Ericson and Wollin, 1968). More recently, the coiling direction of *G.
400 truncatulinoides* has been used to reconstruct paleoceanographic changes in the North
401 Atlantic Ocean since the Early Pleistocene (Billups et al., 2016, 2020; Kaiser et al., 2019).
402 Abundance peaks of sinistral forms of *G. truncatulinoides*, or TE events, have been identified
403 by Ducassou et al. (2018) in sediment cores of the Gulf of Cadiz as remarkable
404 biostratigraphic markers over the last 50 kyrs.

405 At IODP Site U1386, MIS 1 is characterized by a single peak value of *G. truncatulinoides sinistral* ratio. This peak exceeds 87% at 9.1 kyrs cal BP (Fig. 5) and
406 corresponds to TE3 as described by (Ducassou et al., 2018). TE2 and TE1 were not observed
407 in our records, likely due to low sampling resolution and a sedimentary hiatus corresponding
408 to the last ~1.7 kyrs BP.
409

410 During the MIS 12–11 interval, five peaks of *G. truncatulinoides sinistral* ratio
411 exceeding 73% were identified and used to correlate IODP sites U1387 and U1386 (Fig. 6).
412 The top two peaks occur during the transitions between MIS 11b and MIS 11c and MIS 11a
413 and MIS 10. Two peak values exceeding 76% were also observed during MIS 11c. The
414 bottommost "*G.truncatulinoides* event" (TE) with values ranging between 90 and 100 %
415 occurs during MIS 12b. During the MIS 11b, a TE event associated with a contourite level is
416 observed at IODP Site U1386 but the corresponding bioevent and contourite bed are missing
417 at IODP U1387 (Fig. 6). The most likely hypothesis to explain the lack of TE event and of
418 contourite bed in sediment record from IODP U1387 during this period is a short coring gap
419 between cores 339-U1387B-10 and 339-U1387B-11. Alternatively, both TE events and
associated contourite beds, observed during the MIS 11c- MIS 11b interval at IODP Site U1386, could be condensed to a single TE event/contourite bed in the sedimentary record from IODP U1387 (Fig. 7). In all cases, this observation highlights an uncertainty about the age model of IODP Site U1387 during MIS 11b, which has been taken into account when interpreting the data.

<figure><img src="image_8.png" /><figcaption>Fig. 5: Biochronology of IODP sites U1386 and U1387 during the MIS 2–1 interval. IODP Site U1386, from left to right: lithological log; radiocarbon dates and relative abundances of selected planktic foraminifera. IODP Site U1387, from left to right: lithological log; relative abundances of selected planktic foraminifera. MIS: Marine Isotope Stage; GTS: % G. truncatulinoides (s) ratio; YD: Younger Dryas; HS1 to HS3: Heinrich Stadials; mbls: meter below sea level; mcd: meter composite depth. TE: G. truncatulinoides sinistral event; dl: deglaciation I or transition between MIS 2 and MIS 1.</figcaption></figure>
<figure><img src="image_9.png" /></figure>

Fig. 6: Biochronology of IODP sites U1386 and U1387 during the MIS 12–11 interval. For IODP Site U1386, from left to right: stratigraphy from Kaboth et al. (2017) with marine isotope events according to Railsback et al. (2015); lithological log; relative abundances of selected planktic foraminifera. For IODP Site U1387, from left to right: lithological log; relative abundances of selected planktic foraminifera. MIS: Marine Isotope Stage; GTS: % G. truncatulinoides (s) ratio; mcd: meter composite depth. dV: deglaciation V or transition between MIS 12 and MIS 11.

## 5. Results

### 5.1. General sedimentological characteristics

The MIS 2–1 and MIS 12–11 intervals comprise several contourite beds characterized by high values of median grain size, sortable silt and percentage of sand (Figs. Fig. 7–8). The majority of contourite beds observed at IODP Site U1386 during the studied periods, are also present at IODP Site U1387, located 4 km further south-east, reflecting a remarkable spatial continuity. The silty to fine sand contouritic beds occurring during Heinrich Stadial 1 (HS1), the Younger Dryas and the Late Holocene correspond, respectively, to peaks I, II and III
described by Faugères et al. (1986) in the Faro Drift (Fig. 7). During the beginning of deglaciation V (Terminal Stadial Event 12.2; Hodell et al., 2015), the MIS 11c, the transition from substage MIS 11c to MIS 11b, MIS 11b, the transition from the substage MIS 11b to MIS 11a, and at the end of MIS 11a, contourite layers are fine-grained (muddy to silty contourites). However, the analysis of grain-size characteristics indicates that sediments in IODP Site U1387 are coarser than in IODP Site U1386, especially during MIS 11 and MIS 1. During MIS 12.2 and HS1, coarse quartz are present in sediments of IODP Site U1386 but not in sediments of IODP Site U1387 (Figs. 7–8).

Changes in carbonate content during the MIS 12–11 interval echo the Fe/Ca record (Fig. 8). Binocular and microscope observations indicate that carbonates particles are almost exclusively biogenic. The observed similarity between carbonate content and Fe/Ca ratio was used to estimate carbonate content from XRF-derived elemental ratios in other studied intervals. We therefore propose that the Fe/Ca ratio reflects changes in the relative contribution of biogenic vs. terrigenous components. Focussing on glacials, MIS 2 (CaCO₃ ~ 25–30%) is more carbonate-rich than MIS 12 (CaCO₃ ~ 15–25%). Carbonate contents during MIS 11 and MIS 1 are higher (25–45%) than during glacials except during the Early Holocene (20–25%). Located farther from the coast, IODP Site U1387 shows overall higher carbonate contents than IODP Site U1386 during both, MIS 12–11 and MIS 2–1 (Figs. 7–8).
<figure><img src="image_10.png" /><figcaption>Fig. 7: Sedimentological features of IODP sites U1386 and U1387 during MIS 2/1. IODP Site U1386, from left to right: lithological log resulting from visual description with contourite peaks (III, II and I) described by Faugères et al. (1986); grain size characteristics of bulk (black line) and decarbonated (dotted line) sediment: median grain size (D₅₀); sortable silt; percentage of sand; quartz >150 µm; ln(Zr/Al) and ln(Fe/Ca) records. IODP Site U1387, from left to right: lithological log; median grain size (D₅₀); sortable silt; percentage of sand; number of quartz grains (> 150 µm/g dry sediment; ln(Zr/Al) and ln(Fe/Ca) records. HS: Heinrich Stadials; YD: Younger Dryas; MIS: Marine Isotope Stage. III, II and I correspond to contourite peaks described by Faugères et al. (1986); dl: deglaciation I or transition between MIS 2 and MIS 1.</figcaption></figure>

467

468

469

470

471

472

473

474

475
<figure><img src="image_11.png" /><figcaption>Fig. 8: Sedimentological features of IODP sites U1386 and U1387 during MIS 12/11. For both sites from left to right: lithological log resulting from visual description with the identified contourite peaks 1–6; median grain size (D₅₀); sortable silt; percentage of sand; quartz > 150 µm; ln(Zr/Al) and ln(Fe/Ca) records. Black dots: %CaCO₃. 1–6 contourite peaks of MIS 12–11. MIS: Marine Isotope Stage. dV: deglaciation V or transition between MIS 12 and MIS 11. Grey rectangles: G. truncatulinoides events (TE).</figcaption></figure>

## 5.2. Sedimentary facies

Our detailed analysis of sedimentological features in IODP sites U1386 and U1387 complements the lithologies established during IODP Expedition 339 from smear slides (Stow et al., 2013), by revealing the occurrence of 4 main facies and 10 subfacies (Fig. 9).

Facies A is represented by metric intervals of homogenous muddy sediment, and displays the finest grain-size measured in our records with proportions of clays, silts and sands ranging from 75 to 90%, 9 to 24%, and less than 2%, respectively (Fig.9). The D₅₀ is less than 5 µm with a unimodal distribution centered around 3–4 µm.
Facies B forms decimetric to multimetric intervals of silty-clay sediment and is the most abundant facies encountered in our records. The proportion of clays, silts and sands ranges from 44 to 80%, 18 to 51%, and 0 to 12%, respectively (Fig. 7), while the $D_{50}$ spans from 5 to 10 µm. Four subfacies within Facies B were identified according to grain-size distribution and sediment composition. Subfacies B1 shows a bimodal grain-size distribution with a principal mode ranging from 3 to 8 µm and a second mode ranging from 26 to 38 µm. Two compositional subgroups B1a (carbonate-rich) and B1b (terrigenous-rich) can be differentiated within subfacies B1 (Fig. 9). Subfacies B2 displays a unimodal grain size distribution and can be divided in two subgroups due to differences in modal curves and composition of sediment. B2a presents a unimodal modal curve centered around 8–12 µm with a variable proportion of carbonates (20–35%; Fig. 9). With coarser and carbonate-rich sediment (30–35%), B2b is characterized by a unimodal curve centered around 15–22 and slightly fine skewness (Fig. 9).
<table><caption>Fig. 9: Summary of sedimentary facies identified in IODP sites U1386 and U1387. Yellow: CaCO₃ content > 30%; Grey: CaCO₃ content < 30%. *: carbonate content of Facies B2b can be slightly more carbonate rich (35%)</caption><thead><tr><th>Facies</th><th>Grain size</th><th>Modal Curves</th><th>Sand (%)</th><th>Silt (%)</th><th>Clays (%)</th><th>Mode (µm)</th><th>Skewness</th><th>Kurtosis</th><th>General composition (%CaCO3)</th><th>Complete lithologic name (Stow et al., 2012)</th></tr></thead><tbody><tr><td>Facies A</td><td>Muddy</td><td></td><td>0-2</td><td>10-24</td><td>75-90</td><td>Unimodal :4</td><td>Low skewness</td><td>Platycurtic/ Mesocurtic</td><td>Terrigenous (15-30%)</td><td>Calcareous mud</td></tr><tr><td>Facies B1a</td><td rowspan="4">Silty-clay</td><td></td><td>2-5</td><td>24-39</td><td>57-75</td><td>Principal mode :3-6<br>Secondary mode 38-45</td><td>Coarse tail skewed</td><td>Platycurtic</td><td>Terrigenous (15-25%)</td><td>Nannofossils ooze with mud</td></tr><tr><td>Facies B1b</td><td></td><td>0-4</td><td>18-30</td><td>67-80</td><td>Principal mode :3<br>Secondary mode :26-35</td><td>Coarse tail skewed</td><td>Platycurtic/ Mesocurtic</td><td>Carbonate rich (30-40%)</td><td rowspan="2">Nannofossils/calcareous silty mud</td></tr><tr><td>Facies B2a</td><td></td><td>1-5</td><td>33-44</td><td>52-62</td><td>Unimodal :7-12</td><td>Symmetrical</td><td>Platycurtic</td><td>Terrogenous (20-30%) *</td></tr><tr><td>Facies B2b</td><td></td><td>2-12</td><td>37-51</td><td>44-55</td><td>Unimodal :15-22</td><td>Slightly fine skewed</td><td>Platycurtic</td><td>Carbonate rich (30-35%)</td></tr><tr><td>Facies B3</td><td rowspan="2">Silty-mud</td><td></td><td>2-12</td><td>23-43</td><td>50-75</td><td>Bimodal : 3-6 / 26-45</td><td>Symmetrical</td><td>Platycurtic</td><td>Terrigenous (20-30%)</td><td rowspan="2">Silty-mud with biogenic carbonates</td></tr><tr><td>Facies B4</td><td></td><td>9-12</td><td>32-43</td><td>45-59</td><td>Bimodal : 3-6 / 48-65</td><td>Symmetrical</td><td>Platycurtic</td><td>Carbonate rich (> 30%)</td></tr><tr><td>Facies Ca</td><td>Clayey coarse-silt</td><td></td><td>11-27</td><td>29-56</td><td>26-46</td><td>Principal mode :45-78<br>Secondary mode :3-7</td><td>Fine skewed</td><td>Mesocurtic</td><td>Terrigenous (20-30%)</td><td>Sandy-mud with biogenic carbonates</td></tr><tr><td>Facies Cb</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>Carbonate rich (30-35%)</td><td>Silty-muddy calcareous ooze</td></tr><tr><td>Facies D</td><td>Silty-sand</td><td></td><td>31-45</td><td>38-44</td><td>16-25</td><td>Principal mode :65-78</td><td>Very fine skewed</td><td>Leptocurtic</td><td>Carbonate rich (>35%)</td><td>Sandy-silt with biogenic carbonates</td></tr></tbody></table>

505

506

507
508

509

Subfacies B3 presents a bimodal distribution with two modes of grain size ranging from 4 to 8 µm and from 26 to 46 µm, and is principally composed of terrigenous sediment (Fig. 9). Subfacies B4 is characterized by a bimodal distribution with two well separated equal modes ranging from 3 to 6 µm and from 48 to 65 µm, and is more carbonate-rich (Fig. 9). This subfacies is the coarsest B facies with a significant proportion of sand (9–12%) and silt (32–43%). This facies can appear as the “mottled” facies described by Gonthier et al. (1984) in the Gulf of Cadíz. Facies C consists of decimetric intervals of clayey-coarse silt sediment, and is composed of 29 to 46% of clays, 27 to 56% of silts and 11 to 27% of sands.

510

511

512

513

514

515

516

517

518

519

520

521

Its D₅₀ varies between 15 and 21 µm. It presents a bimodal distribution with a dominant coarse mode ranging from 46 to 78 µm and a secondary fine mode ranging from 3 to 9 µm. This facies is also characterized by a mesokurtic grain-size distribution and a fine-tail skewness (Fig. 9). The composition of sediment allows to differentiate two subgroups: a terrigenous subfacies (Ca) and a carbonate-rich subfacies (Cb) (Fig. 9).

522

523

524

525

526

527

Facies D is composed of very thin (centimetric) carbonate-rich silty-sand beds representing the coarsest sediments observed in our archives. The proportion of clays, silts and sands ranges from 16 to 25%, 38 to 44% and 31 to 45%, respectively. D₅₀ ranges from 49 to 57 µm and the grain-size distribution shows a principal mode between 66 and 78 µm. This facies is also characterized by leptokurtic grain-size distribution and a very fine-tail skewness (Fig. 9).

528

529

530
531

5.3.Sedimentary

sequences

of

contourite

beds

532

<figure><img src="image_12.png" /><figcaption>Fig. 10: Detailed sedimentological description of sedimentary sequences 1, 2 (a and b) and 3 V. A, B (1,2,3 and 4), C, D represent sedimentary facies described in the main text. Yellow: $CaCO_3 > 30\%$; Grey: $CaCO_3 < 30\%$.</figcaption></figure>

533

534

535

536

537

538

539

540

541

The gradual facies succession during the periods studied at IODP sites U1386 and U1387 allowed the definition of three sequences (Fig. 10 and Table 4). They are composed of coarsening-up and a fining-up sub-sequences except Holocene sequences which present only a coarsening-up sequence (Sequence 2b, Fig. 10). The thicknesses of coarsening-up and fining-up sub-sequences varies between 63 and 299 cm, and 15 and 203 cm, respectively. We interpret these vertical sequences and their associated grain-size variations as an increase and
542 then a decrease of bottom-current velocity. At this location in the Gulf of Cadiz, beds with
543 such sequences are considered to be contourite layers formed by the MOWU (Faugères et al.,
544 1984; Gonthier et al., 1984).

<table><caption>Table 4: Summary table of the sequences observed in our records. Normal cruces: at IODP Site U1386. Underlined cruces: at IODP Site U1387. MIS: Marine Isotope Stage; YD: Younger Dryas; HS: Heinrich Stadial; dV: deglaciation V or transition between MIS 12 and MIS 11.</caption><thead><tr><th rowspan="2">Sequence</th><th rowspan="2">Facies succession</th><th colspan="5">U1386 - U1387</th><th colspan="2">Interpretation</th></tr><tr><th>Holocene</th><th>MIS 11</th><th>YD</th><th>HS 1</th><th>dV</th><th>Current speed maximum</th><th>Sediment supply trend</th></tr></thead><tbody><tr><td rowspan="2">1</td><td>A-B3-B1b</td><td></td><td></td><td></td><td></td><td>X - X</td><td rowspan="2">Weak (&lt; 10 cm.s<sup>-1</sup>)</td><td>Terrigenous</td></tr><tr><td>B1a-B4-B1a</td><td></td><td>X - X</td><td></td><td></td><td></td><td>Carbonate-rich</td></tr><tr><td rowspan="3">2</td><td>B2a/B2b-B3-Ca-B3-B1b</td><td></td><td></td><td></td><td>X - X</td><td></td><td rowspan="3">Medium (10-15cm.s<sup>-1</sup>)</td><td>Terrigenous</td></tr><tr><td>B1a-B4-Ca-B1a</td><td></td><td>X - X</td><td></td><td></td><td></td><td>Change from carbonate-rich to terrigenous</td></tr><tr><td>B1b-B4-Cb</td><td>X</td><td></td><td></td><td></td><td></td><td>Carbonate-rich</td></tr><tr><td rowspan="2">3</td><td>B2a-Cb-D</td><td>X</td><td></td><td></td><td></td><td></td><td rowspan="2">Strong (15-25cm.s<sup>-1</sup>)</td><td>Carbonate-rich</td></tr><tr><td>B1b-B3-Ca-D-B3-B1b</td><td></td><td></td><td>X - X</td><td></td><td></td><td>Terrigenous</td></tr></tbody></table>

545

546

547

548

549

550

Sequence 1 is the finest sequence encountered in our records and is characteristic of muddy contouritic beds. The increase within this sequence of percentage of silts, its zero or low skewness, its symmetrical distribution of the coarsest facies (B) and the progressive
contact between facies are compatible with very limited bottom current increase (< 10 cm s⁻¹)
(Pepple, 2020). We discriminate two types of muddy contourite beds composing Sequence 1
according to their composition and grain-size characteristics: a terrigenous one (A-B3-B1b)
during dV and a carbonate-rich one (B1a-B4-B1a) during MIS 11 (Fig. 8 and Table 4).

Sequence 2 is a medium grain size sequence typical of silty contouritic beds, and
corresponds to the “Sequence 1” described by Toucanne et al. (2007) during Dansgaard-
Oeschger events in sediments cores of the Gulf of Cadiz beneath the path of MOWU. The
dominance of silts, a fine skewness, and a mesocurtic distribution of the coarsest facies (C)
within this sequence suggest, according to Brackenridge et al. (2018) deposit induced
principally by a weak current (10–15 cm s⁻¹). In addition, the sharp contact between Facies B
and C indicates that Sequence 2 recorded a higher current speed than Sequence 1. Three types
of silty contourite beds composing Sequence 2 were recognized based on their composition
and grain-size characteristics: (1) terrigenous beds (B2-B3-Ca-B3-B1b) observed during HS1
, and (2) Holocene carbonate-rich beds (B1b-B4-Cb) (Fig. 8 and Table 4), both sequences
showing synchronous increases in granularity and carbonate content (Fig. 10); the third type
of silty contouritic beds, more atypical, presents (3) a sequence of carbonate-rich coarsening-
up trend, followed by a more terrigenous maximum of grain-size, and ending with carbonate-
rich fining-up sediments (B1a-B4-Ca-B1a; Fig. 10 and Table 4). This last special sequence is
observed during MIS 11 at both U1386 and U1387 (Table 4).

Sequence 3 is the coarsest sequence encountered in our records and is characteristic of
fine sand contouritic beds. It presents strong similarities with “Sequence 2” proposed by
Toucanne et al. (2007) during the Younger Dryas within a sediment core under the influence
the MOWU. A similar proportion of sand and silt, a very fine skewness and a leptokurtic
distribution of the coarsest facies (D) within this sequence indicate, as proposed by
(Brackenridge et al., 2018), a deposit affected by progressive winnowing at higher current speeds (15–25 cm.s⁻¹). We observe two types of fine sand contouritic beds based on their composition and the contact between individual facies: (1) terrigenous beds (B1b-B3-Ca-D-B3-B1b), characterized by an erosional contact between the Facies Ca and D, and observed during the YD (Fig. 10); (2) Holocene carbonate-rich beds (B2a-Cb-D) observed only at IODP Site U1387 (Table 4) and which present a sharp contact between facies Cb and D. Both types of contourite beds present simultaneous increases and decreases of granularity and carbonate content.

583

## 6.Discussion

585

6.1.Understanding grain size information: current velocity and sediment supply over the Faro Drift

588

6.1.1(bottom current indicators and MOW velocity in the Gulf of Cadíz

590

A majority of paleoceanographical studies in the Gulf of Cadiz used the weight percentage of sands (>63 µm) as an indicator of the intensity of the MOW flow strength (e.g. Rogerson et al., 2005; Llave et al., 2006; Voelker et al., 2006; Kaboth et al., 2017). MOW velocity has also been estimated using the mean of the fine grain size fraction (< 63µm; Voelker et al., 2006, 2015), the 90th centile (D₉₀; Hanquiez et al., 2007), the sorted silt (Toucanne et al., 2007), D90 or by using the median grain-size (Toucanne et al., 2007; Ducassou et al., 2018). Based on a geochemical approach Bahr et al. (2014) showed that the ln(Zr/Al) ratio reflects
bottom current velocity in the Gulf of Cadiz, establishing this ratio as an indicator of MOW
flow strength (Bahr et al., 2015; Voelker et al., 2015). More recently, Brackenridge et al.
(2018) and Pepple (2020) used different grain-size parameters as the median, the sorting, and
the skewness and kurtosis to reconstruct bottom current speed in the Gulf of Cadiz.

In addition to current velocity, other factors such as changes in sediment supply or
depositional processes such as turbidity currents can influence sediment texture (Stow et al.,
1986). Among these factors, the biasing effect of sediment supply changes upon proxy
records of bottom current velocity is barely considered. The sources of sediments to the Faro
Drift are mainly carbonate skeletons of primary and secondary producers, and terrigenous
material delivered by South-West Iberian rivers (Alonso et al., 2016). In the following
discussion we compare the facies succession established in this study with different bottom
current indicator signals in order to have a better understanding of their meanings in the Faro
Drift.

611
<figure><img src="image_13.png" /></figure>

Fig. 11: Close-up of the most recent contourite beds of the Faro Drift. From left to right: contourite beds (peaks) III, II and 4. From top to bottom: (A) ln(Zr/Al) ratio; (B) % of sand fraction (particles > 63 µm) in bulk (black symbols) and carbonate free sediment (dark grey symbols); (C) weight % of the > 63 µm (black symbols) and 63–150 µm fractions (dark grey symbols; Kaboth et al. (2017) and abundance of > 150 µm planktic foraminiferal tests (light grey); (D) sortable silt (dark grey symbols) and D₅₀ (black symbols); (E) facies succession and associated current speed.

We find that the sortable silt is sensitive to slight variations of bottom current velocity of less than 10 cm s⁻¹ (Facies B3 and B4; Brackenridge et al., 2018). However, this grain-size parameter can present identical values for Facies B and C or/and C and D (Fig. 11), thus

612

613

614

615

616

617

618

619

620
demonstrating that the sortable silt can be limited in the accurate reconstruction of bottom current exceeding 10–15 cm s⁻¹. In contrast, the D₅₀ does not always clearly react with minor increases of MOWU velocity (Facies B3 and B4) but instead reliable records medium current speeds (> 10–15 cm s⁻¹).

To reduce the bias caused by changes in sediment supply (foraminiferal or IRD abundances and downslope supply) on the velocity information inferred by the weight percentage of sands, some authors generally take into account the 63–150 µm fraction only (Rogerson et al., 2005; Kaboth et al., 2016, 2017, Sierro et al., 2020). Our results show that sand contributions derived from wet sieving (weight % sand) display significant discrepancies from laser granulometry-derived values (Fig. 11): weight % sand underestimates the amount of sand in the sediment, especially when the sand is terrigenous and mica-rich. Moreover, the weight % sand (> 63 µm and 63–150 µm) present more similarities with downcore profiles of planktic foraminiferal abundance than with gradual facies succession (Fig. 11). These observations suggest that the weight of the coarse fraction reflects both changes in bottom current velocity and changes in the flux of calcareous shells to bottom sediments (Fig. 11).

Representing the relative enrichment of heavy minerals over less dense aluminosilicates under the influence of bottom current flow, the Zr/Al ratio (Bahr et al., 2014) provides a high resolution record of bottom current velocity. This elemental ratio presents similar values between silty-clay facies (B3) of HS1 and YD and clayey coarse-silt facies (Cb) at IODP Site U1386 (Fig. 11). Both facies are principally controlled by the maximum carrying capacity of the current but Facies B3 is normally influenced by a weaker current velocity than facies Cb (Brackenridge et al., 2018). Similarly, at IODP Site U1387 the coarse-silt facies (Ca) of HS1 and YD has ln(Zr/Al) values comparable to those of the silty-sand facies (D) of the Holocene. However, the facies D deposit is affected by non-deposition or progressive winnowing at
higher current speeds (15–25 cm s⁻¹; Pepple, 2020), whereas the facies Ca is not impacted by such high current velocities (10–15 cm s⁻¹; Brackenridge et al., 2018). The “coarse” material available being generally more carbonate-rich during the Holocene, the Zr/Al ratio probably slightly underestimate the MOWU speed during this period.

In the absence of evidence of turbidite events in the Faro Drift area during the studied period (e.g. Gonthier et al., 1984; Hernández-Molina et al., 2003), we relate the vertical distributions of contourite facies at IODP sites U1386 and U1387 during the MIS 12–11 and MIS 2–1 intervals to variations in bottom current velocity.

653

654 6.1.2. *Impact of changes in sediment supply: example of a contouritic sequence of the MIS 11*

655

Unlike most of contourites described in the Faro Drift during the MIS 1 (e.g., Gonthier et al., 1984; Stow et al., 1986) the coarsest part of the MIS 11 contourite sequences does not correspond to maxima in carbonate content (Fig. 11). While the increase of current allows the deposition of carbonate silty particles (facies B4) which is mainly controlled by the carrying capacity of the current, the considerable increase of terrigenous particles in the study area suggests a change in sediment supply. This assumed change in terrigenous sediment delivery to Faro Drift during MIS 11 is concomitant with an increase of *N. pachyderma* reflecting the decrease of sea-surface temperatures (SST) which corresponds to Heinrich-type ice-rafting events (Ht) 2 and 3 detected in Iberian margin records during MIS 11 (Rodrigues et al., 2011; Oliveira et al., 2016). This drop of SST along the Iberian margin has strongly negatively impacted the primary productivity (coccolithophore productivity) and thus the production of secondary producers alike planktic foraminifera during Ht2 and Ht3 (Cavaleiro et al., 2020). The terrigenous biomarkers from southwest Iberian margin indicate, in turn, a
very slight decline (Rodrigues et al., 2011), during the corresponding relatively dry hinterland conditions (Oliveira et al., 2016). These changes in terrigenous and biogenic sediment input in the Faro Drift coincide with the strongest current velocity recorded in these sequences which probably triggered both the winnowing or the non-deposition of terrigenous fine particles and sand sediment supply by saltation. The decrease of carbonate content during the current velocity maxima indicates however, that the increase in terrigenous sediment supply compared to the carbonate supply, such as medium $MOW_U$ velocities, has an influence on sedimentation. Consequently, changes in sediment supply, particularly those induced by changes in productivity vs. export of terrigenous particles can play an important role on the facies succession and contourite sequences in the Gulf of Cadíz (Fig. 11).

## 6.2.Behaviour of the branches of the MOW during MIS 12–11 and MIS 2–1

### 6.2.1.Glacials

The finest facies (A, B1a and B2) identified at IODP sites U1386 and U1387 characterize sediments of MIS 12 and MIS 2 ages. This clay and silty-clay dominated sedimentation as well as the absence of contourite sequences in the studied cores confirm previous studies (Hanquiez, 2006; Llave et al., 2006) which suggest a reduced $MOW_U$ activity during glacials. This reduced activity coincides, during MIS 2, with a deeper circulation of the MOW (e.g., Schönfeld and Zahn, 2000) as well as a strengthening of the lower branch ($MOW_L$). This strengthening is particularly active during HE 2 and HE 3 as shown by the generation of coarse contouritic beds (Llave et al., 2006; Voelker et al., 2006; Toucanne et al., 2007). This general pattern (deeper circulation of the MOW) was probably enhanced during MIS 12
which was characterized by an exceptionally low sea level as well as some of the coldest and driest climate conditions of the Late Pleistocene in the Western Mediterranean region (Sánchez Goñi et al., 2016).

While being both dominated by muddy or silty-clay, U1386 and U1387 sediments of MIS 12 and MIS 2 ages show conspicuous facies discrepancies in-between the two glacials or in-between sites. The MIS 2 facies B2 shows a different modal distribution and a generally higher carbonate concentration than the MIS 12 facies A and B1a. We interpret these differences as an indication of changes in sediment supply between these periods. The extremely low sea-level (-30 m relative to MIS 2) during MIS 12 (Waelbroeck et al., 2002; Rabineau et al., 2006) has contributed to the high export of detrital particles to the middle slope (Mestdagh et al., 2019), and their accumulation in the Faro Drift via their redistribution by the MOW. During MIS 12, the apparent mismatch between records of ln (Fe/Ca) in IODP Site U1386 and biomarkers in a marine core of the southwestern Portuguese margin (MD03-2699), suggests a minor role of coccolithophore productivity and continental inputs (Rodrigues et al., 2011) on the sedimentation in the Faro Drift during this glacial stage.

MIS 12 sediments at IODP sites U1386 and U1387 are characterized by two successive facies, from silty-clay sediment (B1a) at the base of this interval to muddy (A) sediment for the youngest part. This vertical transition to a less coarser facies may reflect a MOWU deepening during MIS 12, and thus a weaker influence of the bottom current on the sedimentation in the Faro Drift.

Interestingly, the carbonate content in sediments of IODP Site U1387 is generally higher than in IODP Site U1386 records. Fine detrital particles, which have mainly a southwestern Iberian origin (Guadalquivir River; Alonso et al., 2016), are redistributed in the Faro Drift by the MOW. Consequently, IODP Site U1387, the furthest site from the core of
717 MOWU, is likely to be less supplied in terrigenous fine particles than IODP Site U1386,
718 explaining the discrepancy in carbonate content between both sites.

719

720 6.2.2.Deglaciations

721

722 Based on the presence of a coarse contouritic bed in the Faro Drift, (“peak I” after
723 Faugères et al. (1986)), many authors argued for an active circulation of the MOW during
724 HS1 (e.g. Mulder et al., 2002; Rogerson et al., 2005; Llave et al., 2006; Voelker et al., 2006;
725 Ducassou et al., 2018). This strengthened flow through the Strait of Gibraltar has been related
726 to the more active formation of WMDW in the Gulf of Lion during Heinrich and stadial
727 events (Rohling et al., 1998; Cacho et al., 2000; 2006; Sierro et al., 2005; Frigola et al., 2008;
728 Bassetti et al, 2010). Our results, in agreement with previous works (Hanquiez, 2006; Llave et
729 al., 2006; Ducassou et al., 2018), show a considerable increase in MOWU activity (Seq2a)
730 which coincides with an enhanced MOWL flow (Mulder et al., 2002; Rogerson et al., 2005;
731 Llave et al., 2006; Voelker et al., 2006; Toucanne et al., 2007; Ducassou et al., 2018).
732 Following Hanquiez (2006), we propose that MOW circulation increased in the Gulf of Cadiz
733 during HS1, and that the lower branch was more active than the upper one. Change in
734 sediment supply plays an important role on the vertical succession of facies during HS1. The
735 youngest part of MIS 2 is characterized by a high contribution of terrigenous particles to
736 sediments of the Faro Drift system. Such an accumulation is due to the high export of detrital
737 particles to the slope during this low sea-level interval, as supported by a stratigraphical study
738 on the continental margin of the northern Gulf of Cadiz (Mestdagh et al., 2019). The
739 proximity of river mouths, as well as the downslope transport in the nepheloid layer and the
740

redistribution by the MOW, may have favoured the high contribution of terrigenous particles
to the sedimentation of the Faro Drift (Stumpf et al., 2011).

741

742

Early on, MOWU strengthening allowed the deposition of terrigenous silty particles (facies B3) controlled principally by the maximum carrying capacity of the current. Then, with a progressive increase in MOWU velocity, sediment accumulation was affected by winnowing. Saltation might also have contributed to the coarse sedimentation recorded at the end of MIS 2, bringing quartz from upstream turbidites deposits during a period of low sea-level (-130/-110 m compared to modern sea-level; Waelbroeck et al., 2002; Elderfield et al., 2012). The rise of carbonate content during this coarsening indicates that winnowing of fine terrigenous particles dominated over the supply of detrital sand by saltation.

743

744

745

746

747

748

749

750

The YD presents the coarsest contourite sequence (Seq 3) identified over the studied periods at IODP sites U1386 and U1387. This sequence contains an erosional contact between facies Ca and D (Fig. 9). Seq3 might be seen as an analogue to coarse contourite beds of YD age recorded in cores under the influence of MOWL (Mulder et al., 2002; Rogerson et al., 2005; Llave et al., 2006; Voelker et al., 2006; Toucanne et al., 2007) and known as "Peak II" (Faugères et al., 1986). We interpret the more carbonate-rich composition of the YD contouritic beds compared to the HS1 beds as induced by a stronger winnowing of the fine terrigenous particles due to high bottom current velocities. A change in terrigenous input has also to be considered since, unlike the HS1 deposits, the YD contourite beds do not contain any coarse quartz (>150 µm). Differences in sea level between these periods (Fig. 12) likely explain the reduced coarse and fine terrigenous particles input in the more distal Faro Drift during the YD. Despite a reduction in detrital sediment supply, the coarseness of facies D as well the presence of an erosional contact within Seq3 point to the highest velocities of the MOWU during the YD (Fig 9). The YD is thus characterized by an enhanced circulation of the

751

752

753

754

755

756

757

758

759

760

761

762

763
lower and upper branches of the MOW as proposed by Hanquiez (2006). The rise in sea level associated with the particularly active circulation of the LIW in the Mediterranean Sea (Toucanne et al., 2012) probably favoured enhanced circulation of the two branches of the MOW during this cold inception (Fig. 12).

The contourite beds (Seq1) observed during the transition from MIS 12 to MIS 11 reflects high activity of the MOWU. Its limited amount of coarse material compared to HS1 (Seq2), a period when sea level was similar to the MIS 12/ MIS 11 transition (Waelbroeck et al., 2002; Elderfield et al., 2012), suggests though that MOWU activity and/or terrigenous supply were lower during dV than during the final part of MIS 2 (Fig. 12).

773

774
<figure><img src="image_14.png" /></figure>

Fig. 12: Upper Mediterranean Outflow Water (MOWU) flow strength and paleoclimatic context. A: Eustatic level after Waelbroeck et al. (2002)(left) and Elderfield et al. (2012) (right). B-C: mean grain-size ($D_{50}$) reflecting MOWU velocity at the sea floor at IODP sites U1386 and U1387. D: mean grain-size ($D_{50}$) of core MD99-2343 collected on the drift of Minorca Islands, approximating Western Mediterranean Deep Water (WMDW) flow speed (Frigola et al., 2008). E: "Sortable Silt" (SS) of core MD01-2472 indicating the Levantine Intermediate Water current intensity along the eastern margin of Corsica (Toucanne et al., 2012). MIS: Marine Isotope Stage; YD: Younger Dryas; HS: Heinrich Stadials; dV(I): deglaciation V(I) or transition between MIS 12(2) and MIS 11(1); V: TE: G. truncatulinoides (s) Events; S: sapropel; Yellow rectangles: contourite beds.

775

776

777

778

779

780

781

782

783
784

785

6.2.3.MIS

11

and

MIS

1

786

787

788

789

790

The presence in our sedimentary records of contourite beds of varying coarseness underscores the enhanced activity of the MOWU during MIS 11 and MIS 1 (Fig. 12) as suggested by Rogerson et al. (2005), Llave et al. (2006), Bahr et al. (2014), and Kaboth et al. (2016).

791

792

793

794

795

796

797

798

799

800

801

802

803

804

805

806

807

During the MIS 11 and Holocene high-sea level periods, contourite beds are coarser and more carbonate-rich at IODP Site U1387 which is located at a greater distance from the moat, than at IODP Site U1386. Two factors can be put forward to explain these spatial compositional and grain-size discrepancies: differences in bottom current speed due to intensification and/or spatial variability of the MOWU, and/or differences in sediment supply. When located higher in the water column, the core of the MOWU has a greater influence at IODP Site U1387 during MIS 11 and MIS 1 because of its slightly shallower position than at IODP Site U1386. When sea level is lower, especially at the beginning of deglaciations (dV, HS1), the core of the current migrates downwards, affecting both U1386 and U1387. During MIS 11 and MIS 1, deposition of fine terrigenous particles in the studied area has been limited by (1) the high sea-level (Waelbroeck et al., 2002) which favoured the deposition of a portion of detrital clays in the continental shelf and prevented the export of terrigenous particles to the middle slope and (2) the active MOWU which limited the deposition of fine particles, in particular during its strengthening. These findings are supported by the overall higher sedimentation rates during glacial intervals than during interglacials in the Faro Drift (Kaboth et al., 2017).
808

809

810

811

812

813

814

815

The “Peak III” coarse contourite bed described by Faugères et al. (1986) and many authors in the Faro Drift (Mulder et al., 2002; Llave et al., 2006; Ducassou et al., 2018) and observed at IODP sites U1386 and U1387, is also recorded in our archives during the Holocene. It reflects a slight intensification of the MOWU around 6 kyrs cal BP and then a considerable strengthening of the flow over the last 3 kyrs (Fig. 12). These accelerations of the MOWU are probably due to the renewed ventilation of the LIW 6 kyrs ago (Toucanne et al., 2012) and a strengthened deep-water formation (WMDW) 4 kyrs ago in the Gulf of Lion (Frigola et al., 2008).

816

817

818

819

820

821

822

823

824

825

Our results show that Peak III dating from the end of Holocene is much coarser than the contouritic beds encountered in MIS 11. The increases of MOWU flow strength during the transition from the interglacial (MIS 11c) to the following glacial (MIS 10), took place while sea level was more or less similar to the Holocene sea level (Waelbroeck et al., 2002; Elderfield et al., 2012). Moreover, during both periods relatively higher primary productivity and associated secondary production of foraminifera has been observed south-west off the Iberian peninsula (Rodrigues et al., 2009; 2011). Consequently, despite similarities in the sedimentary processes, the difference of coarseness of contourite beds between both MIS 11 and MIS 1 suggests that MOWU velocity was higher during the end of Holocene than during MIS 11 despite no evidence of sequence finning-up.

826

827

828

829

830

831

832

The MIS 11 sedimentary hiatus observed in the MOWL-influenced IODP Site U1390 corresponds to the Late Quaternary Discontinuity in the south Portuguese CDS, which is related to tectonic activity and uplift-induced erosion (Hernández-Molina et al., 2016; Lofi et al., 2016). However, the possibility that this hiatus was also induced by a contemporary strengthening of MOWL can not be ruled out. Indeed, the benthic foraminiferal assemblages described in sediments of IODP Site U1391 off south-west Portugal indicate that MOWL
833 reached its highest intensity of the last 900 ka during MIS 11 (Guo et al., 2017). These results
834 highlight a slightly different flow pattern of the MOW between MIS 11 and MIS 1. While the
835 most recent interglacial is characterized by an intense MOW_U flow (Rogerson et al., 2005;
836 Llave et al., 2006; Bahr et al., 2014; Kaboth et al., 2016) and a less active MOW_L circulation
837 (Mulder et al., 2002; Rogerson et al., 2005; Llave et al. 2006; Voelker et al. 2006; Toucanne
838 et al. 2007), our results indicate that MOW_U circulation during MIS 11 was lower while the
839 MOW_L circulation was particularly intense.

840 A relationship between sapropel formation in the Eastern Mediterranean Basin and low
841 MOW activity was first demonstrated for Sapropel 1 (S1) during the Holocene (Voelker et al.,
842 2006; Toucanne et al., 2007; Rogerson et al., 2012; Kaboth et al., 2016). It was then extended
843 for all sapropels down to MIS 7 (Sierro et al., 2020). Our results confirm the low activity of
844 the MOW_U during S1 and highlight a reduced MOW_U flow during S11 (418.9 kyrs cal. BP;
845 Konijnendijk et al., 2014; Fig. 12).
846

6.3. MOW strengthening during deglaciation V and MIS 11: a strong influence on

847

North Atlantic climate?

<figure><img src="image_15.png" /></figure>
849

850

851

852

853

854

Fig. 13: Mediterranean Upper Water variability and North Atlantic paleoceanography. A) epibenthic δ¹³C record from IODP Site U1313, North Atlantic. B) weight % of IRD (grey), relative abundances of subpolar planktic foraminiferal species (blue Oppo et al., 1998) and benthic δ¹³C record at ODP Site 980, North Atlantic (McManus et al., 1999). C) Benthic foraminiferal δ¹⁸O values (Kaboth et al., 2017) and relative abundances of subpolar planktic foraminifera at IODP Site U1386, Gulf of Cadiz. C: Sortable silt (orange symbols) and mean grain-size (black) at IODP Site U1386, Gulf of Cadiz. MIS: Marine Isotope Stage. Yellow rectangles: contourite beds.

855

856

857

858

859

860

861

862

863

864

865

866

Model simulations suggest that the AMOC is significantly influenced by the contribution of the MOW (Reid, 1979; Rahmstorf, 1998; Wu et al., 2007) but the effects of Mediterranean water masses on AMOC remain unclear (Swingedouw et al., 2018). Indeed, while a strengthening of MOW does not influence North Atlantic Deep Water production (NADW; Khélifi et al., 2014), it stimulates the shallow branch of AMOC (Bahr et al., 2015; Sarnthein et al., 2018). Sediment records have shown a negative feedback between exchanges at the Strait of Gibraltar and AMOC during the Late Pliocene and the Late Quaternary (Rogerson et al., 2012; García-Gallardo et al., 2018). As previously inferred from other marine records in the Gulf of Cadiz, the intensification of the MOW could have played a role in the intensification or reinvigoration of AMOC, during the last deglaciation (Rogerson et al., 2006), at the end of Heinrich Stadials (Voelker et al., 2006) or during MIS5e (Bahr et al., 2015).

867

868

869

870

871

872

873

874

The activation of the MOW<sub>U</sub> (1) is observed just before the massive terminal ice rafting event in the North Atlantic occurring during deglaciation V (Oppo et al., 1998) and (2) is contemporaneous with a decrease of benthic δ¹³C values at North Atlantic sites (ODP Site 980 and IODP Site U1313), which is interpreted as a diminution of North Atlantic Deep Water production (Fig. 13). Despite variations of intensity, the general behaviour of the MOW was probably similar during deglaciation V and the last deglaciation. Consequently, as proposed by many authors for the deglaciation I (e.g., Rogerson et al., 2006; Voelker et al., 2006), MOW may have played an important role in maintaining the AMOC, and particularly
its upper branch (Sarnthein et al., 2018), during the period of weak NADW formation which characterized the deglaciation V. However, the stimulation of the upper AMOC during the Heinrich Stadials by increasing salinity at intermediate depth and the subsequent transport of moisture to the high latitudes (Bahr et al., 2015), had the potential to strongly impact global climate. This supply of moisture may have contributed to the growth of especially large northern hemisphere ice sheets which may have, in turn, contributed to their destabilization. Thus participating to the “excess” of ice on ice sheets at the end of MIS 12, which was the major factor of Heinrich Stadial of deglaciation V (Vázquez Riveiros et al., 2013; Bahr et al., 2018; Catunda et al., 2021).

MIS 11c is, in our study area, characterized by a weak MOWU activity. The slight coarsening of sortable silt associated with the D50 which remains stable, only reflects a very low increase of the bottom current velocity (Fig. 13). Simultaneously, the increase of benthic δ13C values at North Atlantic records, indicates a strong NADW production during the MIS 11c (McManus et al., 1999; Voelker et al., 2010). By contrast, at the transition between the MIS 11c and MIS 11b, there is a significant strengthening of MOWU reflected by the simultaneous increase of sortable silt and D50 and a progressive reduction of NADW formation (Fig. 13). These results support previous suggestions of a negative feedback between deep AMOC during MIS 11. This increase of MOWU activity at the boundary between MIS 11c and MIS 11b (Fig. 13) may have also strengthened the upper AMOC but the diminished size and stability of the ice sheet during MIS 11c (McManus et al., 1999) may explain the lack of related IRD in the North Atlantic (Oppo et al., 1998). However, we observe a concomitant increase in both, the abundance of *N. pachyderma* at ODP Site 980, indicative of a drop of SST in North Atlantic, and the flow of the MOWU (Fig. 13). The moisture supplied by the MOW to high latitude at the end of interglacial MIS 11c may have
favour the growth of ice-sheet, and thus played a role in the transition to the substage MIS 11b.

The behaviour of the MOWU is similar between the beginning of MIS 11c and the Holocene. During the early part of each interglacial, reduced MOWU is observed, associated with a sapropel formation in the Eastern Mediterranean Sea. Conversely, the end of each period is characterized by pulses of the MOWU, but with a considerable difference of bottom current velocity between both interglacials. Consequently, although our results support that North Atlantic overall circulation is similar during MIS 11c and the Holocene (McManus et al., 1999; Rodrigues et al., 2011), it should not be neglected that MOWU, current intensities might have differed. Interestingly, assuming that the end of the current interglacial would occur in the near future (1.5 kyrs) with CO2 concentrations lower than 240 ppmv (Tzedakis et al., 2012a), MOW pulses of MIS 11 and Holocene would occur few thousand years (3–6 kyrs) before the onset of the glaciation.

During MIS 11b and at the end of MIS 11a, MOWU strengthening preceded cold events (stadials) in the North Atlantic Ocean as reflected by IRD discharges at ODP Site 980 (Oppo et al., 1998) (Fig. 13). As proposed by Bahr et al. (2015) for MIS 5e, pulses of MOW can invigorate the upper AMOC by increasing the salinity of North Atlantic Ocean water masses at intermediate depth. Such a stimulated upper branch of the overturning circulation contributed to heat and moisture transport towards the high latitudes during this time, thus sustaining initial ice-sheet growth (McManus et al., 1999). The recorded high abundances of *N. pachyderma* at ODP Site 980, which correspond to slight increases of the percentages of the polar planktic species at IODP sites U1386 and U1387 during MIS 11a, indicate that the slight increases of calving or/and size of ice sheets caused a general cooling in the North Atlantic (Fig. 13). SST drops during the MIS 11 were also recorded in the mid-latitudes of the
924 North Atlantic Ocean (U1313) and to the southwest of the Iberian margin (Voelker et al., 2010; Rodrigues et al., 2011). The significant increases of MOWU activity during MIS 11b 925 and at the end of MIS 11a correspond to slight decreases of benthic δ¹³C values in North 926 Atlantic, which indicate short episodes of decreased NADW production (Fig. 13). These 927 observations further confirm the negative feedback between the deep branch of the AMOC 928 and exchanges of water masses at the Strait of Gibraltar during the MIS 11. These results, as 929 those of García-Gallardo et al. (2018) for the late Pliocene, support the existence of a 930 “negative feedback” between Mediterranean and North Atlantic deep circulation (Rogerson et 931 al., 2010). 932

933

934 7.Conclusions

935

936 Our records from IODP sites U1386 and U1387 show that the composition of 937 contourite sequences in the Faro Drift, Gulf of Cadíz, is constrained by a balance between 938 MOWU velocity and sediment supplies. Using sedimentary facies as an indicator of MOW 939 strength, our data extend the model of MOWU circulation, previously established for the last 940 140 kyrs, to the MIS 12–11 periods, showing an active MOWU during warm intervals and 941 deglaciations. 942

942 Our grain-size and compositional data obtained at IODP sites U1386 and U1387 show 943 clear evidence that, being more influenced by river plumes during periods of lower sea level, 944 contourite sequences have a more terrigenous composition during deglaciations than during 945 interglacials. During deglaciations, in addition to the increase of the bottom current velocity, 946 this input of terrigenous particles may have played an important role on contourite sequences. 947
Our detailed study of MIS 11 contourite sequences revealed compositional and grain size peculiarities, which are unique to Faro Drift contourite beds in the Gulf of Cadíz. Contrary to the general pattern displayed by contourite beds in this area, grain-size maxima are not associated with peaks in carbonate content but with higher contributions of terrigenous particles. Rapid changes of environmental conditions in the Gulf of Cadíz leading to sudden increases of river discharges and drastic reductions of foraminiferal populations in the surface waters are thought to explain these intra MIS 11 variabilities in contourite sequences.

Our unprecedented, high resolution record of MOWU intensity during MIS 12–11 indicates that strengthened circulation of this water mass during deglaciation V and MIS 11 may have had a significant impact on upper AMOC revigurations and associated heat and moisture transport toward the high latitudes.

## Acknowledgements

P. Moal-Darrigade is supported by a doctoral scholarship of the French ministry of research. The authors thank the crew and the scientific teams of the IODP Expedition 339 on the *Joides Resolution*. We acknowledge the help provided by the Bremen Core Repository staff during the sampling periods. We also wish to thank J. Niard, E. Brung, A. Chabrerie and L. Devaux for their assistance during sample preparation. Finally, the authors thank the French programmes INSU LEFE (IMAGO) for financial support and INSU Artémis for most of the radiocarbon dating of this study. We also thank Antje Voelker, for his constructive review, which help us to improve the final version of the manuscript.

## References
969 Alonso, B., Ercilla, G., Casas, D., Stow, D.A.V., Rodríguez-Tovar, F.J., Dorador, J.,

970 Hernández-Molina, F.-J., 2016. Contourite vs gravity-flow deposits of the Pleistocene

971 Faro Drift (Gulf of Cadiz): Sedimentological and mineralogical approaches. Marine

972 Geology 377, 77–94. https://doi.org/10.1016/j.margeo.2015.12.016

973 Ambar, I., 1983. A shallow core of Mediterranean water off western Portugal. Deep Sea

974 Research Part A. Oceanographic Research Papers 30, 677–680.

975 Ambar, I., Howe, M., 1979. Observations of the Mediterranean outflow—I mixing in the

976 Mediterranean outflow. Deep Sea Research Part A. Oceanographic Research Papers

977 26, 535–554.

978 Ambar, I., Serra, N., Brogueira, M.J., Cabeçadas, G., Abrantes, F., Freitas, P., Gonçalves, C.,

979 Gonzalez, N., 2002. Physical, chemical and sedimentological aspects of the

980 Mediterranean outflow off Iberia. Deep Sea Research Part II: Topical Studies in

981 Oceanography 49, 4163–4177. https://doi.org/10.1016/S0967-0645(02)00148-0

982 Bahr, A., Jiménez-Espejo, F.J., Kolasinac, N., Grunert, P., Hernández-Molina, F.J., Röhl, U.,

983 Voelker, A.H.L., Escutia, C., Stow, D.A.V., Hodell, D., Alvarez-Zarikian, C.A., 2014.

984 Deciphering bottom current velocity and paleoclimate signals from contourite deposits

985 in the Gulf of Cádiz during the last 140 kyr: An inorganic geochemical approach.

986 Geochem. Geophys. Geosyst. 15, 3145–3160. https://doi.org/10.1002/2014GC005356

987 Bahr, A., Kaboth, S., Hodell, D., Zeeden, C., Fiebig, J., Friedrich, O., 2018. Oceanic heat

988 pulses fueling moisture transport towards continental Europe across the mid-

989 Pleistocene transition. Quaternary Science Reviews 179, 48–58.

990 https://doi.org/10.1016/j.quascirev.2017.11.009

991 Bahr, A., Kaboth, S., Jiménez-Espejo, F.J., Sierro, F.J., Voelker, A.H.L., Lourens, L., Röhl,

992 U., Reichart, G.J., Escutia, C., Hernández-Molina, F.J., Pross, J., Friedrich, O., 2015.
993 Persistent monsoonal forcing of Mediterranean Outflow Water dynamics during the

994 late Pleistocene. Geology 43, 951–954. https://doi.org/10.1130/G37013.1

995 Baringer, M.O., Price, J.F., 1999. A review of the physical oceanography of the

996 Mediterranean outflow. Marine Geology 155, 63–82. https://doi.org/10.1016/S0025-

997 3227(98)00141-8

998 Be, A.W.H., 1960. Ecology of Recent planktonic foraminifera--Part 2, Bathymetric and

999 seasonal distributions in the Sargasso Sea off Nermuda. Micropaleontology 6, 373–

1000 392.

1001 Bigg, G.R., Wadley, M.R., 2001. Millennial-scale variability in the oceans: an ocean

1002 modelling view. Journal of Quaternary Science 16, 309–319.

1003 https://doi.org/10.1002/jqs.599

1004 Billups, K., Hudson, C., Kunz, H., Rew, I., 2016. Exploring Globorotalia truncatulinoides

1005 coiling ratios as a proxy for subtropical gyre dynamics in the northwestern Atlantic

1006 Ocean during late Pleistocene Ice Ages. Paleoceanography 31, 553–563.

1007 https://doi.org/10.1002/2016PA002927

1008 Billups, K., Vizcaíno, M., Chiarello, J., Kaiser, E.A., 2020. Reconstructing Western Boundary

1009 Current Stability in the North Atlantic Ocean for the Past 700 Kyr From Globorotalia

1010 truncatulinoides Coiling Ratios. Paleoceanography and Paleoclimatology 35,

1011 e2020PA003958. https://doi.org/10.1029/2020PA003958

1012 Borenäs, K., Wåhlin, A., Ambar, I., Serra, N., 2002. The Mediterranean outflow splitting—a

1013 comparison between theoretical models and CANIGO data. Deep Sea Research Part

1014 II: Topical Studies in Oceanography 49, 4195–4205.

1015 Bowen, D.Q., 2010. Sea level ~400 000 years ago (MIS 11): analogue for present and future

1016 sea-level? Climate of the Past 6, 19–29. https://doi.org/10.5194/cp-6-19-2010
1017 Brackenridge, R.E., Hernández-Molina, F.J., Stow, D.A.V., Llave, E., 2013. A Pliocene
mixed contourite–turbidite system offshore the Algarve Margin, Gulf of Cadiz:
1018
Seismic response, margin evolution and reservoir implications. Marine and Petroleum
1019
Geology 46, 36–50. https://doi.org/10.1016/j.marpetgeo.2013.05.015
1020

1021 Brackenridge, R.E., Stow, D.A.V., Hernández-Molina, F.J., Jones, C., Mena, A., Alejo, I.,
1022 Ducassou, E., Llave, E., Ercilla, G., Nombela, M.A., Perez-Arlucea, M., Frances, G.,
1023 2018. Textural characteristics and facies of sand-rich contourite depositional systems.
1024 Sedimentology 65, 2223–2252. https://doi.org/10.1111/sed.12463

1025 Brunner, C.A., 1986. Deposition of a muddy sediment drift in the southern Straits of Florida
during the late Quaternary. Marine Geology 69, 235–249.
1026
1027 https://doi.org/10.1016/0025-3227(86)90041-1

1028 Bryden, H.L., Candela, J., Kinder, T.H., 1994. Exchange through the Strait of Gibraltar.
1029 Progress in Oceanography 33, 201–248. https://doi.org/10.1016/0079-6611(94)90028-0
1030 0

1031 Bryden, H.L., Stommel, H.M., 1984. Limiting processes. Oceanol. Acta 7, 289–296.

1032 Carracedo, L.I., Pardo, P.C., Flecha, S., Pérez, F.F., 2016. On the Mediterranean Water
1033 Composition. Journal of Physical Oceanography 46, 1339–1358.
1034 https://doi.org/10.1175/JPO-D-15-0095.1

1035 Catunda, M.C.A., Bahr, A., Kaboth-Bahr, S., Zhang, X., Foukal, N.P., Friedrich, O., 2021.
1036 Subsurface Heat Channel Drove Sea Surface Warming in the High-Latitude North
1037 Atlantic During the Mid-Pleistocene Transition. Geophysical Research Letters 48,
e2020GL091899. https://doi.org/10.1029/2020GL091899

1039 Cavaleiro, C., Voelker, A.H.L., Stoll, H., Baumann, K.-H., Kucera, M., 2020.

1040 Coccolithophore productivity at the western Iberian Margin during the Middle
1041 Pleistocene (310-455&thinsp;ka) – evidence from coccolith Sr/Ca data. Climate of the
1042 Past 16, 2017-2037. https://doi.org/10.5194/cp-16-2017-2020

1043 de Abreu, L., Abrantes, F.F., Shackleton, N.J., Tzedakis, P.C., McManus, J.F., Oppo, D.W.,
1044 Hall, M.A., 2005. Ocean climate variability in the eastern North Atlantic during
1045 interglacial marine isotope stage 11: A partial analogue to the Holocene?
1046 Paleoceanography 20. https://doi.org/10.1029/2004PA001091

1047 Demirov, E.K., Pinardi, N., 2007. On the relationship between the water mass pathways and
1048 eddy variability in the Western Mediterranean Sea. J. Geophys. Res. 112.
1049 https://doi.org/10.1029/2005JC003174

1050 Droxler, A.W., Farrell, J.W., 2000. Marine Isotope Stage 11 (MIS 11): new insights for a
1051 warm future. Global and Planetary Change 24, 1-5. https://doi.org/10.1016/S0921-8181(99)00065-X
1052

1053 Ducassou, E., Hassan, R., Gonthier, E., Duprat, J., Hanquiez, V., Mulder, T., 2018.
1054 Biostratigraphy of the last 50 kyr in the contourite depositional system of the Gulf of
1055 Cádiz. Marine Geology 395, 285-300. https://doi.org/10.1016/j.margeo.2017.09.014

1056 Elderfield, H., Ferretti, P., Greaves, M., Crowhurst, S., McCave, I.N., Hodell, D., Piotrowski,
1057 A.M., 2012. Evolution of Ocean Temperature and Ice Volume Through the Mid-
1058 Pleistocene Climate Transition. Science 337, 704-709.
1059 https://doi.org/10.1126/science.1221294

1060 Ericson, D.B., Wollin, G., 1968. Pleistocene Climates and Chronology in Deep-Sea
1061 Sediments. Science 162, 1227-1234.

1062 Expedition 339 Scientists, 2012. Mediterranean outflow: environmental significance of the
1063 Mediterranean Outflow Water and its global implications. IODP Prel. Rept., 339.

1064 Eynaud, F., de Abreu, L., Voelker, A., Schönfeld, J., Salgueiro, E., Turon, J.-L., Penaud, A.,
1065 Toucanne, S., Naughton, F., Sánchez Goñi, M.F., Malaizé, B., Cacho, I., 2009.
1066 Position of the Polar Front along the western Iberian margin during key cold episodes
1067 of the last 45 ka. Geochemistry, Geophysics, Geosystems 10.
1068 https://doi.org/10.1029/2009GC002398
1069 Faugeres, J.-C., Gonthier, E., Peypouquet, J., Pujol, C., Vergnaud-Grazzini, C., 1986.
1070 Distribution et variations des courants de fond sur la ride de Faro (Golfe de Cadix),
1071 témoins des modifications des échanges Méditerranée-Atlantique au Quaternaire
1072 récent. Bulletin de la Société Géologique de France 2, 423-432.
1073 Faugères, J.-C., Gonthier, E., Stow, D.A.V., 1984. Contourite drift molded by deep
1074 Mediterranean outflow. Geology 12, 296-300. https://doi.org/10.1130/0091-
1075 7613(1984)12<296:CDMBDM>2.0.CO;2
1076 Frigola, J., Moreno, A., Cacho, I., Canals, M., Sierro, F.J., Flores, J.A., Grimalt, J.O., 2008.
1077 Evidence of abrupt changes in Western Mediterranean Deep Water circulation during
1078 the last 50kyr: A high-resolution marine record from the Balearic Sea. Quaternary
1079 International 181, 88-104. https://doi.org/10.1016/j.quaint.2007.06.016
1080 García, M., Hernández-Molina, F.J., Llave, E., Stow, D.A.V., León, R., Fernández-Puga,
1081 M.C., Diaz del Río, V., Somoza, L., 2009. Contourite erosive features caused by the
1082 Mediterranean Outflow Water in the Gulf of Cadiz: Quaternary tectonic and
1083 oceanographic implications. Marine Geology 257, 24-40.
1084 https://doi.org/10.1016/j.margeo.2008.10.009
1085 García-Gallardo, Á., Grunert, P., Piller, W.E., 2018. Variations in Mediterranean-Atlantic
1086 exchange across the late Pliocene climate transition. Clim. Past 14, 339-350.
1087 https://doi.org/10.5194/cp-14-339-2018
1088 Gonthier, E., Faugères, J.-C., Stow, D., 1984. Contourite facies of the Faro drift, Gulf of
1089 Cadiz. Geological Society, London, Special Publications 15, 275-292.
1090 https://doi.org/10.1144/GSL.SP.1984.015.01.18
1091 Gonthier, E.G., Faugères, J.-C., Stow, D.A.V., 1984. Contourite facies of the Faro Drift, Gulf of Cadiz. Geological Society, London, Special Publications 15, 275-292.

1092 https://doi.org/10.1144/GSL.SP.1984.015.01.18

1094 Hanquiez, V., Mulder, T., Lecroart, P., Gonthier, E., Marchès, E., Voisset, M., 2007. High resolution seafloor images in the Gulf of Cadiz, Iberian margin. Marine Geology 246, 42-59. https://doi.org/10.1016/j.margeo.2007.08.002

1097 Harris, P.T., Brancolini, G., Armand, L., Busetti, M., Beaman, R.J., Giorgetti, G., Presti, M., Trincardi, F., 2001. Continental shelf drift deposit indicates non-steady state Antarctic bottom water production in the Holocene. Marine Geology 179, 1-8. https://doi.org/10.1016/S0025-3227(01)00183-9

1101 Heezen, B.C., Hollister, C.D., Ruddiman, W.F., 1966. Shaping of the continental rise by deep geostrophic contour currents. Science 152, 502-508.

1103 Hernández-Molina, F.J., Llave, E., Preu, B., Ercilla, G., Fontan, A., Bruno, M., Serra, N., Gomiz, J.J., Brackenridge, R.E., Sierro, F.J., Stow, D.A.V., García, M., Juan, C., Sandoval, N., Arnaiz, A., 2014. Contourite processes associated with the Mediterranean Outflow Water after its exit from the Strait of Gibraltar: Global and conceptual implications. Geology 42, 227-230. https://doi.org/10.1130/G35083.1

1108 Hernández-Molina, F.J., Llave, E., Stow, D.A.V., García, M., Somoza, L., Vázquez, J.T., Lobo, F.J., Maestro, A., Díaz del Río, V., León, R., Medialdea, T., Gardner, J., 2006. The contourite depositional system of the Gulf of Cádiz: A sedimentary model related to the bottom current activity of the Mediterranean outflow water and its interaction with the continental margin. Deep Sea Research Part II: Topical Studies in Oceanography 53, 1420-1463. https://doi.org/10.1016/j.dsr2.2006.04.016
1114 Hernández-Molina, F.J., Serra, N., Stow, D.A.V., Llave, E., Ercilla, G., Van Rooij, D., 2011.

1115 Along-slope oceanographic processes and sedimentary products around the Iberian

1116 margin. Geo-Mar Lett 31, 315–341. https://doi.org/10.1007/s00367-011-0242-2

1117 Hernández-Molina, F.J., Sierro, F.J., Llave, E., Roque, C., Stow, D.A.V., Williams, T., Lofi,

1118 J., Van der Schee, M., Arnáiz, A., Ledesma, S., Rosales, C., Rodríguez-Tovar, F.J.,

1119 Pardo-Igúzquiza, E., Brackenridge, R.E., 2016. Evolution of the gulf of Cadiz margin

1120 and southwest Portugal contourite depositional system: Tectonic, sedimentary and

1121 paleoceanographic implications from IODP expedition 339. Marine Geology 377, 7–

1122 39. https://doi.org/10.1016/j.margeo.2015.09.013

1123 Hernández-Molina, F.J., Stow, D., Alvarez-Zarikian, C., Expedition IODP 339 Scientists,

1124 2013. IODP Expedition 339 in the Gulf of Cadiz and off West Iberia: decoding the

1125 environmental significance of the Mediterranean outflow water and its global

1126 influence. Sci. Dril. 16, 1–11. https://doi.org/10.5194/sd-16-1-2013

1127 Hernández-Molina, J., Llave, E., Somoza, L., Fernández-Puga, M.C., Maestro, A., León, R.,

1128 Medialdea, T., Barnolas, A., García, M., del Río, V.D., Rodero, J., Gardner, J., 2003.

1129 Looking for clues to paleoceanographic imprints: a diagnosis of the Gulf of Cadiz

1130 contourite depositional systems. Geology 31, 19–22. https://doi.org/10.1130/0091-

1131 7613(2003)031<0019:LFCTPI>2.0.CO;2

1132 Hodell, D., Lourens, L., Crowhurst, S., Konijnendijk, T., Tjallingii, R., Jiménez Espejo, F.,

1133 Skinner, L., Tzedakis, P.C., Abrantes, F., Acton, G.D., Alvarez Zarikian, C.A., Bahr,

1134 A., Balestra, B., Llave Barranco, E., Carrara, G., Ducassou, E., Flood, R.D., Flores, J.-

1135 A., Furota, S., Grimalt, J., Grunert, P., Hernández-Molina, J., Kim, J.K., Krissek, L.A.,

1136 Kuroda, J., Li, B., Lofi, J., Margari, V., Martrat, B., Miller, M.D., Nanayama, F.,

1137 Nishida, N., Richter, C., Rodrigues, T., Rodríguez-Tovar, F.J., Freixo Roque, A.C.,

1138 Sánchez Goñi, M.F., Sierro Sánchez, F.J., Singh, A.D., Sloss, C.R., Stow, D.A.V.,
1139 Takashimizu, Y., Tzanova, A., Voelker, A., Xuan, C., Williams, T., 2015. A reference time scale for Site U1385 (Shackleton site) on the SW Iberian Margin. Affiliation (analytic): University of Cambridge, Department of Earth Sciences, Cambridge
1141 Affiliation (monographic): University of Cambridge, Department of Earth Sciences, Cambridge, United Kingdom Corporate Affiliation (monographic): Shackleton Site
1142 Project Members Coordinates: N373417 N373417 W0100719 W0100720 illus., incl.
1143 table, sketch map Contains 61 references Research Program: IODP Integrated Ocean
1145 Drilling Program Document Type: Journal Article Bibliographic Level: Analytic
1146 Source Note: Global and Planetary Change, Vol.133, p.49-64. Publisher: Elsevier,
1148 Amsterdam, Netherlands. ISSN: 0921-8181 Copyright Information: GeoRef,
1149 Copyright 2019 American Geosciences Institute. Reference includes data from
1150 CAPCAS, Elsevier Scientific Publishers, Amsterdam, Netherlands GeoRef ID:
1151 2016037553 DOI: 10.1016/j.gloplacha.2015.07.002 133, 49.

1152 Ivanovic, R.F., Valdes, P.J., Flecker, R., Gregoire, L.J., Gutjahr, M., 2013. The parameterisation of Mediterranean-Atlantic water exchange in the Hadley Centre model HadCM3, and its effect on modelled North Atlantic climate. Ocean Modelling 62, 11-16. https://doi.org/10.1016/j.ocemod.2012.11.002

1156 Jouzel, J., Masson-Delmotte, V., Cattani, O., Dreyfus, G., Falourd, S., Hoffmann, G., Minster, B., Nouet, J., Barnola, J.M., Chappellaz, J., Fischer, H., Gallet, J.C., Johnsen, S., Leuenberger, M., Loulergue, L., Luethi, D., Oerter, H., Parrenin, F., Raisbeck, G., Raynaud, D., Schilt, A., Schwander, J., Selmo, E., Souchez, R., Spahni, R., Stauffer, B., Steffensen, J.P., Stenni, B., Stocker, T.F., Tison, J.L., Werner, M., Wolff, E.W., 2007. Orbital and Millennial Antarctic Climate Variability over the Past 800,000 Years. Science 317, 793-796. https://doi.org/10.1126/science.1141038
1163 Kaboth, S., Bahr, A., Reichart, G.-J., Jacobs, B., Lourens, L.J., 2016. New insights into upper MOW variability over the last 150kyr from IODP 339 Site U1386 in the Gulf of Cadiz. Marine Geology 377, 136–145. https://doi.org/10.1016/j.margeo.2015.08.014
1164
1165
1166 Kaboth, S., de Boer, B., Bahr, A., Zeeden, C., Lourens, L.J., 2017a. Mediterranean Outflow Water dynamics during the past ~570 kyr: Regional and global implications: Middle to Late Pleistocene MOW. Paleoceanography 32, 634–647.
1167
1168
1169 https://doi.org/10.1002/2016PA003063
1170 Kaboth, S., Grunert, P., Lourens, L., 2017b. Mediterranean Outflow Water variability during the Early Pleistocene. Clim. Past 13, 1023–1035. https://doi.org/10.5194/cp-13-1023-2017
1171
1172
1173 Kaiser, E.A., Caldwell, A., Billups, K., 2019. North Atlantic Upper-Ocean Hydrography During the Mid-Pleistocene Transition Evidenced by Globorotalia truncatulinoides Coiling Ratios. Paleoceanography and Paleoclimatology 34, 658–671.
1174
1175
1176 https://doi.org/10.1029/2018PA003502
1177 Khélifi, N., Sarnthein, M., Frank, M., Andersen, N., Garbe-Schönberg, D., 2014. Late Pliocene variations of the Mediterranean outflow. Marine Geology 357, 182–194.
1178
1179 https://doi.org/10.1016/j.margeo.2014.07.006
1180 Konijnendijk, T.Y.M., Ziegler, M., Lourens, L.J., 2014. Chronological constraints on Pleistocene sapropel depositions from high-resolution geochemical records of ODP Sites 967 and 968. Newsletters on Stratigraphy 263–282.
1181
1182
1183 https://doi.org/10.1127/0078-0421/2014/0047
1184 Lisiecki, L.E., Raymo, M.E., 2005. A Pliocene-Pleistocene stack of 57 globally distributed benthic δ¹⁸O records: PLIOCENE-PLEISTOCENE BENTHIC STACK.
1185
1186
1186 Paleoceanography 20, n/a-n/a. https://doi.org/10.1029/2004PA001071
1187 Llave, E., Hernández-Molina, F.J., Somoza, L., Stow, D., Del Río, V.D., 2007. Quaternary evolution of the contourite depositional system in the Gulf of Cadiz. Geological Society, London, Special Publications 276, 49–79.

1190 Llave, E., Schönfeld, J., Hernández-Molina, F.J., Mulder, T., Somoza, L., Díaz del Río, V., Sánchez-Almazo, I., 2006. High-resolution stratigraphy of the Mediterranean outflow contourite system in the Gulf of Cadiz during the late Pleistocene: The impact of Heinrich events. Marine Geology 227, 241–262.
1194 https://doi.org/10.1016/j.margeo.2005.11.015

1195 Lofi, J., Voelker, A.H.L., Ducassou, E., Hernández-Molina, F.J., Sierro, F.J., Bahr, A., Galvani, A., Lourens, L.J., Pardo-Igúzquiza, E., Pezard, P., Rodríguez-Tovar, F.J., Williams, T., 2016. Quaternary chronostratigraphic framework and sedimentary processes for the Gulf of Cadiz and Portuguese Contourite Depositional Systems derived from Natural Gamma Ray records. Marine Geology 377, 40–57.
1200 https://doi.org/10.1016/j.margeo.2015.12.005

1201 Loutre, M.F., Berger, A., 2003. Marine Isotope Stage 11 as an analogue for the present interglacial. Global and Planetary Change 36, 209–217.
1203 https://doi.org/10.1016/S0921-8181(02)00186-8

1204 Loutre, M.F., Berger, A., 2000. Future Climatic Changes: Are We Entering an Exceptionally Long Interglacial? Climatic Change 46, 61–90.
1206 https://doi.org/10.1023/A:1005559827189

1207 Madelain, F., 1970. Influence de la topographie du fond sur l'écoulement méditerranéen entre le Détroit de Gibraltar et le Cap Saint-Vincent. Cahiers océanographiques 22, 43–61.

1209 Marchès, E., Mulder, T., Cremer, M., Bonnel, C., Hanquiez, V., Gonthier, E., Lecroart, P., 2007. Contourite drift construction influenced by capture of Mediterranean Outflow Water deep-sea current by the Portimão submarine canyon (Gulf of Cadiz, South
1212 Portugal). Marine Geology 242, 247–260.

1213 https://doi.org/10.1016/j.margeo.2007.03.013

1214 McCave, I.N., Hall, I.R., 2006. Size sorting in marine muds: Processes, pitfalls, and prospects

1215 for paleoflow-speed proxies: SIZE SORTING IN MARINE MUDS. Geochemistry,

1216 Geophysics, Geosystems 7, n/a-n/a. https://doi.org/10.1029/2006GC001284

1217 McCave, I.N., Manighetti, B., Robinson, S.G., 1995. Sortable silt and fine sediment

1218 size/composition slicing: Parameters for palaeocurrent speed and palaeoceanography.

1219 Palaeoceanography 10, 18. https://doi.org/10.1029/94PA03039

1220 McManus, J., Oppo, D., Cullen, J., Healey, S., 2003. Marine Isotope Stage 11 (Mis 11):

1221 Analog for Holocene and Future Climate?, in: Earth’s Climate and Orbital

1222 Eccentricity: The Marine Isotope Stage 11 Question. American Geophysical Union

1223 (AGU), pp. 69–85. https://doi.org/10.1029/137GM06

1224 McManus, J.F., Oppo, D.W., Cullen, J.L., 1999. A 0.5-Million-Year Record of Millennial-

1225 Scale Climate Variability in the North Atlantic. Science 283, 971–975.

1226 https://doi.org/10.1126/science.283.5404.971

1227 Mestdagh, T., Lobo, F.J., Llave, E., Hernández-Molina, F.J., Van Rooij, D., 2019. Review of

1228 the late Quaternary stratigraphy of the northern Gulf of Cadiz continental margin:

1229 New insights into controlling factors and global implications. Earth-Science Reviews

1230 198, 102944. https://doi.org/10.1016/j.earscirev.2019.102944

1231 Millot, C., 2014. Heterogeneities of in-and out-flows in the Mediterranean Sea. Progress in

1232 oceanography 120, 254–278. https://doi.org/10.1016/j-pocean.2013.09.007

1233 Millot, C., 2009. Another description of the Mediterranean Sea outflow. Progress in

1234 Oceanography 82, 101–124. https://doi.org/10.1016/j-pocean.2009.04.016
1235 Millot, C., Candela, J., Fuda, J.-L., Tber, Y., 2006. Large warming and salinification of the Mediterranean outflow due to changes in its composition. Deep Sea Research Part I: Oceanographic Research Papers 53, 656-666.

1238 Mulder, T., Hassan, R., Ducassou, E., Zaragosi, S., Gonthier, E., Hanquiez, V., Marchès, E., Toucanne, S., 2013. Contourites in the Gulf of Cadiz: a cautionary note on potentially ambiguous indicators of bottom current velocity. Geo-Mar Lett 33, 357-367.
1241 https://doi.org/10.1007/s00367-013-0332-4

1242 Mulder, T., Lecroart, P., Hanquiez, V., Marches, E., Gonthier, E., Guedes, J.-C., Thiébot, E., Jaaidi, B., Kenyon, N., Voisset, M., Perez, C., Sayago, M., Fuchey, Y., Bujan, S., 2006. The western part of the Gulf of Cadiz: contour currents and turbidity currents interactions. Geo-Mar Lett 26, 31-41. https://doi.org/10.1007/s00367-005-0013-z

1246 Naranjo, C., Sammartino, S., García-Lafuente, J., Bellanco, M.J., Taupier-Letage, I., 2015. Mediterranean waters along and across the Strait of Gibraltar, characterization and zonal modification. Deep Sea Research Part I: Oceanographic Research Papers 105, 41-52. https://doi.org/10.1016/j.dsr.2015.08.003

1250 Nelson, C.H., Maldonado, A., 1999. The Cadiz margin study off Spain: an introduction. Marine Geology 155, 3-8. https://doi.org/10.1016/S0025-3227(98)00138-8

1252 Oliveira, D., Desprat, S., Rodrigues, T., Naughton, F., Hodell, D., Trigo, R., Rufino, M., Lopes, C., Abrantes, F., Goni, M.F.S., 2016. The complexity of millennial-scale variability in southwestern Europe during MIS 11. Quat. res. 86, 373-387.
1255 https://doi.org/10.1016/j.yqres.2016.09.002

1256 Oppo, D.W., McManus, J.L., Cullen, J.L., 1998. Abrupt Climate Events 500,000 to 340,000 Years Ago: Evidence from Subpolar North Atlantic Sediments. Science 279, 1335-1338. https://doi.org/10.1126/science.279.5355.1335
1259 Palumbo, E., Voelker, A.H.L., Flores, J.A., Amore, O.F., 2019. Surface-ocean dynamics during eccentricity minima: a comparison between interglacial Marine Isotope Stage (MIS) 1 and MIS 11 on the Iberian Margin. Global and Planetary Change 172, 242–255. https://doi.org/10.1016/j.gloplacha.2018.10.011

1263 Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.-M., Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kotlyakov, V.M., Legrand, M., Lipenkov, V.Y., Lorius, C., PÉpin, L., Ritz, C., Saltzman, E., Stievenard, M., 1999. Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica. Nature 399, 429–436. https://doi.org/10.1038/20859

1268 Rabineau, M., Berné, S., Olivet, J.-L., Aslanian, D., Guillocheau, F., Joseph, P., 2006. Paleo sea levels reconsidered from direct observation of paleoshoreline position during Glacial Maxima (for the last 500,000 yr). Earth and Planetary Science Letters 252, 119–137. https://doi.org/10.1016/j.epsl.2006.09.033

1272 Rahmstorf, S., 1998. Influence of mediterranean outflow on climate. Eos Trans. AGU 79, 281–281. https://doi.org/10.1029/98EO00208

1274 Railsback, L.B., Gibbard, P.L., Head, M.J., Voarintsoa, N.R.G., Toucanne, S., 2015. An optimized scheme of lettered marine isotope substages for the last 1.0 million years, and the climatostratigraphic nature of isotope stages and substages. Quaternary Science Reviews 111, 94–106. https://doi.org/10.1016/j.quascirev.2015.01.012

1278 Raynaud, D., Barnola, J.-M., Souchez, R., Lorrain, R., Petit, J.-R., Duval, P., Lipenkov, V.Y., 2005. The record for marine isotopic stage 11. Nature 436, 39–40. https://doi.org/10.1038/43639b

1281 Rebesco, M., Camerlenghi, A., Loon, A.J.V., 2008. Contourite Research: a field in full development. Developements in Sedimentology 60, 0–8.
1283 Rebesco, M., Hernández-Molina, F.J., Van Rooij, D., Wåhlin, A., 2014. Contourites and associated sediments controlled by deep-water circulation processes: State-of-the-art and future considerations. Marine Geology 352, 111–154.
https://doi.org/10.1016/j.margeo.2014.03.011

1287 Reid, J.L., 1979. On the contribution of the Mediterranean Sea outflow to the Norwegian-Greenland Sea. Deep Sea Research Part A. Oceanographic Research Papers 26, 1199–1223.
https://doi.org/10.1016/0198-0149(79)90064-5

1290 Reimer, P.J., Baillie, M.G.L., Bard, E., Bayliss, A., Beck, J.W., Blackwell, P.G., Ramsey, C.B., Buck, C.E., Burr, G.S., Edwards, R.L., Friedrich, M., Grootes, P.M., Guilderson, T.P., Hajdas, I., Heaton, T.J., Hogg, A.G., Hughen, K.A., Kaiser, K.F., Kromer, B., McCormac, F.G., Manning, S.W., Reimer, R.W., Richards, D.A., Southon, J.R., Talamo, S., Turney, C.S.M., 2009. IntCal09 and Marine09 radiocarbon age calibration curves, 0–50,000 years CAL BP. Radiocarbon 51, 1111–1150.

1296 Robinson, S.G., McCave, I.N., 1994. Orbital forcing of bottom-current enhanced sedimentation on Feni Drift, NE Atlantic, during the mid-Pleistocene.
Paleoceanography 9, 943–972.
https://doi.org/10.1029/94PA01439

1299 Rodrigues, T., Grimalt, J.O., Abrantes, F.G., Flores, J.A., Lebreiro, S.M., 2009. Holocene interdependences of changes in sea surface temperature, productivity, and fluvial inputs in the Iberian continental shelf (Tagus mud patch). Geochem. Geophys. Geosyst. 10, n/a-n/a.
https://doi.org/10.1029/2008GC002367

1303 Rodrigues, T., Voelker, A.H.L., Grimalt, J.O., Abrantes, F., Naughton, F., 2011. Iberian Margin sea surface temperature during MIS 15 to 9 (580–300 ka): Glacial suborbital variability versus interglacial stability.
Paleoceanography 26, PA1204.
https://doi.org/10.1029/2010PA001927
1307 Rogerson, M., Colmenero-Hidalgo, E., Levine, R.C., Rohling, E.J., Voelker, A.H.L., Bigg,

1308 G.R., Schönfeld, J., Cacho, I., Sierro, F.J., Löwemark, L., Reguera, M.I., de Abreu, L.,

1309 Garrick, K., 2010. Enhanced Mediterranean-Atlantic exchange during Atlantic

1310 freshening phases. Geochem. Geophys. Geosyst. 11, n/a-n/a.

1311 https://doi.org/10.1029/2009GC002931

1312 Rogerson, M., Rohling, E.J., Bigg, G.R., Ramirez, J., 2012. Paleoceanography of the Atlantic-

1313 Mediterranean exchange: Overview and first quantitative assessment of climatic

1314 forcing. Rev. Geophys. 50. https://doi.org/10.1029/2011RG000376

1315 Rogerson, M., Rohling, E.J., Weaver, P.P.E., 2006. Promotion of meridional overturning by

1316 Mediterranean-derived salt during the last deglaciation. Paleoceanography 21.

1317 https://doi.org/10.1029/2006PA001306

1318 Rogerson, M., Rohling, E.J., Weaver, P.P.E., Murray, J.W., 2005. Glacial to interglacial

1319 changes in the settling depth of the Mediterranean Outflow plume. Paleoceanography

1320 20. https://doi.org/10.1029/2004PA001106

1321 Rohling, E.J., Braun, K., Grant, K., Kucera, M., Roberts, A.P., Siddall, M., Trommer, G.,

1322 2010. Comparison between Holocene and Marine Isotope Stage-11 sea-level histories.

1323 Earth and Planetary Science Letters 291, 97–105.

1324 https://doi.org/10.1016/j.epsl.2009.12.054

1325 Sanchez Goñi, M.F., Harrison, S.P., 2010. Millennial-scale climate variability and vegetation

1326 changes during the Last Glacial: Concepts and terminology. Quaternary Science

1327 Reviews 29, 2823–2827. https://doi.org/10.1016/j.quascirev.2009.11.014

1328 Sánchez Goñi, M.F., Llave, E., Oliveira, D., Naughton, F., Desprat, S., Ducassou, E., Hodell,

1329 D.A., Hernández-Molina, F.J., 2016. Climate changes in south western Iberia and

1330 Mediterranean Outflow variations during two contrasting cycles of the last 1Myrs:
1331 MIS 31–MIS 30 and MIS 12–MIS 11. Global and Planetary Change 136, 18–29.

1332 https://doi.org/10.1016/j.gloplacha.2015.11.006

1333 Sarnthein, M., Grunert, P., Khélifi, N., Frank, M., Nürnberg, D., 2018. Interhemispheric

1334 teleconnections: Late Pliocene change in Mediterranean outflow water linked to

1335 changes in Indonesian Through-Flow and Atlantic Meridional Overturning

1336 Circulation, a review and update. Int J Earth Sci (Geol Rundsch) 107, 505–515.

1337 https://doi.org/10.1007/s00531-017-1505-6

1338 Schiebel, R., Hemleben, C., 2017. Classification and Taxonomy of Extant Planktic

1339 Foraminifers, in: Schiebel, R., Hemleben, C. (Eds.), Planktic Foraminifers in the

1340 Modern Ocean. Springer, Berlin, Heidelberg, pp. 11–110. https://doi.org/10.1007/978-

1341 3-662-50297-6_2

1342 Schönfeld, J., Zahn, R., 2000. Late Glacial to Holocene history of the Mediterranean Outflow.

1343 Evidence from benthic foraminiferal assemblages and stable isotopes at the

1344 Portuguese margin. Palaeogeography, Palaeoclimatology, Palaeoecology 159, 85–111.

1345 https://doi.org/10.1016/S0031-0182(00)00035-3

1346 Schroeder, K., Chiggiato, J., Josey, S.A., Borghini, M., Aracri, S., Sparnocchia, S., 2017.

1347 Rapid response to climate change in a marginal sea. Scientific Reports 7, 1–7.

1348 Serra, N., Ambar, I., Boutov, D., 2010. Surface expression of Mediterranean Water dipoles

1349 and their contribution to the shelf/slope – open ocean exchange. Ocean Sci. 19.

1350 Siegenthaler, U., Stocker, T.F., Monnin, E., Lüthi, D., Schwander, J., Stauffer, B., Raynaud,

1351 D., Barnola, J.-M., Fischer, H., Masson-Delmotte, V., Jouzel, J., 2005. Stable Carbon

1352 Cycle–Climate Relationship During the Late Pleistocene. Science 310, 1313–1317.

1353 https://doi.org/10.1126/science.1120130
1354 Sierro, F.J., Flores, J.A., Baraza, J., 1999. Late glacial to recent paleoenvironmental changes in the Gulf of Cadiz and formation of sandy contourite layers. Marine Geology 155, 157-172. https://doi.org/10.1016/S0025-3227(98)00145-5

1357 Sierro, F.J., Hodell, D.A., Andersen, N., Azibeiro, L.A., Jimenez-Espejo, F.J., Bahr, A., Flores, J.A., Ausin, B., Rogerson, M., Lozano-Luz, R., others, 2020. Mediterranean overflow over the last 250 kyr: freshwater forcing from the tropics to the ice sheets. Paleooceanography and Paleoclimatology 35, e2020PA003931. https://doi.org/10.1029/2020PA003931

1362 Spahni, R., Chappellaz, J., Stocker, T.F., Loulergue, L., Hausmann, G., Kawamura, K., Flückiger, J., Schwander, J., Raynaud, D., Masson-Delmotte, V., Jouzel, J., 2005.

1364 Atmospheric Methane and Nitrous Oxide of the Late Pleistocene from Antarctic Ice Cores. Science 310, 1317-1321. https://doi.org/10.1126/science.1120132

1366 Spezzaferri, S., Kucera, M., Pearson, P.N., Wade, B.S., Rappo, S., Poole, C.R., Morard, R., Stalder, C., 2015. Fossil and Genetic Evidence for the Polyphyletic Nature of the Planktonic Foraminifera "Globigerinoides", and Description of the New Genus Trilobatus. PLoS ONE 10, e0128108. https://doi.org/10.1371/journal.pone.0128108

1370 Stow, D.A., Faugères, J.-C., Howe, J.A., Pudsey, C.J., Viana, A.R., 2002. Bottom currents, contourites and deep-sea sediment drifts: current state-of-the-art. Geological Society, London, Memoirs 22, 7-20.

1373 Stow, D.A., Hernández-Molina, F.J., Llave, E., Sayago-Gil, M., Díaz del Río, V., Branson, A., 2009. Bedform-velocity matrix: the estimation of bottom current velocity from bedform observations. Geology 37, 327-330.

1376 Stow, D.A.V., Faugères, J.-C., Gonthier, E., 1986a. Facies distribution and textural variation in Faro Drift contourites: Velocity fluctuation and drift growth. Marine Geology 72, 71-100. https://doi.org/10.1016/0025-3227(86)90100-3
1379 Stow, D.A.V., Faugères, J.-C., Gonthier, E., 1986b. Facies distribution and textural variation in Faro Drift contourites: Velocity fluctuation and drift growth. Marine Geology 72, 71–100. https://doi.org/10.1016/0025-3227(86)90100-3

1382 Stow, D.A.V., Hernández-Molina, F.J., Llave, E., Bruno, M., García, M., Díaz del Rio, V., Somoza, L., Brackenridge, R.E., 2013. The Cadiz Contourite Channel: Sandy contourites, bedforms and dynamic current interaction. Marine Geology 343, 99–114. https://doi.org/10.1016/j.margeo.2013.06.013

1386 Stumpf, R., Frank, M., Schönfeld, J., Haley, B.A., 2011. Climatically driven changes in sediment supply on the SW Iberian shelf since the Last Glacial Maximum. Earth and Planetary Science Letters 312, 80–90. https://doi.org/10.1016/j.epsl.2011.10.002

1389 Swingedouw, D., Colin, C., Eynaud, F., Ayache, M., Zaragosi, S., 2019. Impact of freshwater release in the Mediterranean Sea on the North Atlantic climate. Clim Dyn 53, 3893–3915. https://doi.org/10.1007/s00382-019-04758-5

1392 Telford, R.J., Heegaard, E., Birks, H.J.B., 2004. The intercept is a poor estimate of a calibrated radiocarbon age. The Holocene 14, 296–298. https://doi.org/10.1191/0959683604h1707fa

1395 Tjallingii, R., Röhl, U., Kölling, M., Bickert, T., 2007. Influence of the water content on X-ray fluorescence core-scanning measurements in soft marine sediments. Geochemistry, Geophysics, Geosystems 8. https://doi.org/10.1029/2006GC001393

1398 Toucanne, S., Jouet, G., Ducassou, E., Bassetti, M.-A., Dennielou, B., Angue Minto'o, C.M., Lahmi, M., Touyet, N., Charlier, K., Lericolais, G., Mulder, T., 2012. A 130,000-year record of Levantine Intermediate Water flow variability in the Corsica Trough, western Mediterranean Sea. Quaternary Science Reviews 33, 55–73. https://doi.org/10.1016/j.quascirev.2011.11.020
1403 Toucanne, S., Mulder, T., Schönfeld, J., Hanquiez, V., Gonthier, E., Duprat, J., Cremer, M.,

1404 Zaragosi, S., 2007. Contourites of the Gulf of Cadiz: A high-resolution record of the

1405 paleocirculation of the Mediterranean outflow water during the last 50,000 years.

1406 Palaeogeography, Palaeoclimatology, Palaeoecology 246, 354–366.

1407 https://doi.org/10.1016/j.palaeo.2006.10.007

1408 Turon, J.-L., Lézine, A.-M., Denèfle, M., 2003. Land–sea correlations for the last glaciation

1409 inferred from a pollen and dinocyst record from the Portuguese margin. Quat. res. 59,

1410 88–96. https://doi.org/10.1016/S0033-5894(02)00018-2

1411 Tzedakis, P.C., Channell, J.E.T., Hodell, D.A., Kleiven, H.F., Skinner, L.C., 2012a.

1412 Determining the natural length of the current interglacial. Nature Geosci 5, 138–141.

1413 https://doi.org/10.1038/ngeo1358

1414 Tzedakis, P.C., Wolff, E.W., Skinner, L.C., Brovkin, V., Hodell, D.A., McManus, J.F.,

1415 Raynaud, D., 2012b. Can we predict the duration of an interglacial? Climate of the

1416 Past 8, 1473–1485. https://doi.org/10.5194/cp-8-1473-2012

1417 Vázquez Riveiros, N., Waelbroeck, C., Skinner, L., Duplessy, J.-C., McManus, J.F.,

1418 Kandiano, E.S., Bauch, H.A., 2013. The “MIS 11 paradox” and ocean circulation:

1419 Role of millennial scale events. Earth and Planetary Science Letters 371–372, 258–

1420 268. https://doi.org/10.1016/j.epsl.2013.03.036

1421 Voelker, A., Lebreiro, S., Schönfeld, J., Cacho, I., Erlenkeuser, H., Abrantes, F., 2006.

1422 Mediterranean outflow strengthening during northern hemisphere coolings: A salt

1423 source for the glacial Atlantic? Earth and Planetary Science Letters 245, 39–55.

1424 https://doi.org/10.1016/j.epsl.2006.03.014

1425 Voelker, A.H.L., de Abreu, L., Schönfeld, J., Erlenkeuser, H., Abrantes, F., 2009.

1426 Hydrographic conditions along the western Iberian margin during marine isotope stage

1427 2. Geochem. Geophys. Geosyst. 10. https://doi.org/10.1029/2009GC002605
1428 Voelker, A.H.L., Rodrigues, T., Billups, K., Oppo, D., McManus, J., Stein, R., Hefter, J., Grimalt, J.O., 2010. Variations in mid-latitude North Atlantic surface water properties during the mid-Brunhes (MIS 9–14) and their implications for the thermohaline circulation. Clim. Past 6, 531–552. https://doi.org/10.5194/cp-6-531-2010

1432 Voelker, A. H. L., Salgueiro, E., Rodrigues, T., Jimenez-Espejo, F.J., Bahr, A., Alberto, A., Loureiro, I., Padilha, M., Rebotim, A., Roehl, U., 2015. Mediterranean Outflow and surface water variability off southern Portugal during the early Pleistocene: A snapshot at Marine Isotope Stages 29 to 34 (1020–1135 ka). Global and Planetary Change 133, 223–237. https://doi.org/10.1016/j.gloplacha.2015.08.015

1437 Voelker, Antje H.L., Salgueiro, E., Rodrigues, T., Jimenez-Espejo, F.J., Bahr, A., Alberto, A., Loureiro, I., Padilha, M., Rebotim, A., Röhl, U., 2015. Mediterranean Outflow and surface water variability off southern Portugal during the early Pleistocene: A snapshot at Marine Isotope Stages 29 to 34 (1020–1135 ka). Global and Planetary Change 133, 223–237. https://doi.org/10.1016/j.gloplacha.2015.08.015

1442 Waelbroeck, C., Labeyrie, L., Michel, E., Duplessy, J.C., McManus, J.F., Lambeck, K., Balbon, E., Labracherie, M., 2002. Sea-level and deep water temperature changes derived from benthic foraminifera isotopic records. Quaternary Science Reviews 21, 295–305. https://doi.org/10.1016/S0277-3791(01)00101-9

1446 Wu, W., Danabasoglu, G., Large, W.G., 2007. On the effects of parameterized Mediterranean overflow on North Atlantic ocean circulation and climate. Ocean Modelling 19, 31–52. https://doi.org/10.1016/j.ocemod.2007.06.003

1449

1450