# Changes in the deep-water masses of the Eastern Mediterranean Sea during the last interglacial maximum - New constraints from neodymium isotopes in foraminifera

Guohui Gao, Christophe Colin, Giuseppe Siani, Sophie Sepulcre, Zhifei Liu, Jiawang Wu, Arnaud Dapoigny, Elisabeth Michel, Rosella Pinna, Amélie Plautre, et al.

► To cite this version:

Guohui Gao, Christophe Colin, Giuseppe Siani, Sophie Sepulcre, Zhifei Liu, et al. Changes in the deep-water masses of the Eastern Mediterranean Sea during the last interglacial maximum - New constraints from neodymium isotopes in foraminifera. Quaternary Science Reviews, 2025, 360, pp.109391. 10.1016/j.quascirev.2025.109391. hal-05093854

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# Changes in the deep-water masses of the Eastern Mediterranean Sea during the last interglacial maximum - New constraints from neodymium isotopes in foraminifera

Guohui Gao<fnref n="1" />, Christophe Colin<fnref n="1" />*, Giuseppe Siani<fnref n="1" />, Sophie Sepulcre<fnref n="1" />, Zhifei Liu<fnref n="2" />, Jiawang Wu<fnref n="3" />,

Arnaud Dapoigny<fnref n="4" />, Elisabeth Michel<fnref n="4" />, Rosella Pinna<fnref n="1" />, Amélie Plautre<fnref n="1" />, Nejib Kallel<fnref n="5" />, Paolo

Montagna<fnref n="6" />

## Abstract

The thermohaline circulation of the Mediterranean Sea was a key factor in the development of sapropel layers in the Eastern Mediterranean Sea (EMS). Sapropel formation is linked to stagnant deep-water conditions and/or increased sea surface productivity during boreal summer insolation maximum. However, debates persist regarding the complex interplay between high and low latitude climatic processes, circulation dynamics, and sapropel formation. Analyses of major and trace elements by XRF, δ¹⁸O of planktonic foraminifera, and neodymium isotopes (εNd) of foraminifera have been carried out on two cores collected in the southeastern Aegean Sea of the EMS and in the Sardinia Channel of the Western Mediterranean Sea (WMS), in order to constrain

<footnote n="1">¹Université Paris-Saclay, CNRS, GEOPS, Orsay 91405, France</footnote>

<footnote n="2">²State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, China</footnote>

<footnote n="3">³School of Marine Sciences, Sun Yat-Sen University, 519082 Zhuhai, China</footnote>

<footnote n="4">⁴Laboratoire des Sciences du Climat et de L'environnement, LSCE/IPSL, CEA-CNRS-UVSQ, Université Paris-Saclay, Gif-sur-Yvette 91191, France</footnote>

<footnote n="5">⁵Faculty of Sciences of Sfax, GEOGLOB Laboratory, University of Sfax, Sfax, Tunisia</footnote>

<footnote n="6">⁶Institute of Polar Sciences, ISP-CNR, Via Gobetti 101, 40129 Bologna, Italy</footnote>

<footnote>*Corresponding author: christophe.colin@universite-paris-saclay.fr (CC)</footnote>
deep-water mass circulation during the last glacial-interglacial cycle (i.e. the past ~150 ka), focusing on the last interglacial maximum. When combined with previous εNd records from the EMS, we identified two regional patterns of the εNd variations during the sapropel S5 interval (~128-120 ka), suggesting that the initial freshwater influx from North Africa led to marked water stratification, limited deep-water formation, and shifted foraminiferal εNd towards a local signature due to an extended period of Nd exchange between seawater and marginal sediments. Cores strongly influenced from the Nile River discharge display radiogenic εNd values (-3.7), whereas those from the northern EMS exhibit more unradiogenic εNd values (-5.1). As S5 progressed, εNd records of all cores from the Levantine basin indicate a slight mixing of deep-water masses accompanied by a decrease in the εNd gradient between the northern and southern EMS. This was associated with the global decrease in temperature and reduced Nile River water input, which supported a slight deep-water recirculation from the Aegean Sea. This active circulation conveyed radiogenic εNd from the southern EMS to the north and the Western Mediterranean Sea. Overall, our new results suggest that hydrological conditions in Mediterranean Sea fluctuated during the S5 deposition, with a slight recirculation of Aegean deep-water in a pre-stage before the complete recovery of circulation in the EMS at the end of the sapropel S5 deposition.

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**Key words:** Neodymium isotope composition; foraminifera; paleoceanography; sapropels; Mediterranean Sea.

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# 1. Introduction

The sedimentation of the Mediterranean Sea is characterized by the periodic deposition of sapropels, a phenomenon that has been documented since the upper Miocene period (Cramp and O'Sullivan, 1999; De Kaenel et al., 1999; Emeis et al., 1991; Meyers and Doose, 1999; Murat, 1999; Meyers, 2006; Rogerson et al., 2008; Rohling et al., 2015). Sapropels are dark gray to olive green and black marine sediments, characterized by enhanced content of total organic carbon (TOC) than the surrounding sediments (Cita et al., 1977; Calvert, 1983; Kidd et al., 1978; Kullenberg, 1952; Murat and Got, 1987; Olausson, 1961; Ryan and Cita, 1977; Vergnaud-Grazzini
et al., 1977, 1985). The formation of sapropels in the Eastern Mediterranean Sea (EMS) is directly linked to: (1) reduced deep-water ventilation associated with freshwater input lowering surface water salinity that leads to stratification of the water masses, resulting in limited or no oxygen renewal in deep water; or (2) nutrient-rich surface freshwater inputs, which significantly increase surface productivity and induce the remineralization of organic matter in deep water to a level that exceeds oxygen inputs through deep water mass renewal; or (3) a combination of both processes (Cramp and O'Sullivan, 1999; De Lange et al., 2008; Emeis et al., 2000; Rossignol-Strick, 1983; Rohling, 1994; Rohling et al., 2015; Schmiedl et al., 2010; Ziegler et al., 2010; ). However, the complex interactions between hydrological circulation, climate conditions at high and low latitudes, and sapropel deposition continue to be a subject of extensive debate (Colin et al., 2021; Rohling et al., 2015; Wu et al., 2019).

A widely accepted hypothesis to explain the excess freshwater input during times of sapropel formation is that it may be related to heavy monsoonal precipitation in North Africa, with the resulting runoff flowing into the EMS via the Nile and other North African paleo-rivers (e.g., Blanchet et al., 2021; Emeis et al., 2003; Fontugne et al., 1994; Osborne et al., 2008, 2010; Rossignol-Strick et al., 1982; Rohling et al., 2002; Revel et al., 2010; Scrivner et al., 2004; Wu et al., 2016, 2017, 2019). In North Africa, the occurrence of humid periods has been attributed to the northward migration of the rain belt associated with the Inter-Tropical Convergence Zone (ITCZ), as a consequence of precession-driven insolation changes (Arbuszewski et al., 2013; Demenocal et al., 2000; Gasse, 2000; Rossignol-Strick et al., 1982; Skonieczny et al., 2015, 2019). Conversely, the intervals between sapropels are characterized by a reduction in riverine runoff and an increase in wind-blown sediment supplies (e.g., Larrasoana et al., 2003; Revel et al., 2015; Wehausen and Brumsack, 1999; Zhao et al., 2012). However, other studies based on pollen data in Padul lake sediments in Southern Spain (Camuera et al., 2022), pollen and isotope record in Morocco (Cheddadi et al., 2021) and the oxygen isotope gradient within of the North African fossil aquifers (Abouelmagd et al., 2012; Edmund et al., 2003) show that humid periods in North Africa and the Western Mediterranean Sea (WMS) were of North Atlantic origin and were linked to the southward shift of the Mediterranean storm track.

North Africa experienced a period of exceptionally high precipitation known as the African
Humid Period 5 (AHP5) from ~132 to ~126 ka during the last interglacial maximum corresponding to the Marine Isotopic Stages (MIS) 5e (Bastian et al., 2017; Ehrmann et al., 2016; Shanahan et al., 2015). This period of intense rainfall was even more pronounced than that of the last African Humid Period (AHP1: from 14.8 to 6 cal ka BP, Bastian et al., 2017; Shanahan et al., 2015). The AHP5 is thought to be responsible for the formation of the organic-rich S5 deposit between 128 and 120 ka in the EMS (Grant et al., 2016), which is one of the most developed sapropels of the late Quaternary, with TOC concentrations typically reaching 7-8% and occasionally exceeding 15% (e.g., Cane et al., 2002; Marino et al., 2009; Osborne et al., 2008, 2010; Rohling et al., 2002, 2006; Struck et al., 2001; Scrivner et al., 2004). On the other hand, several studies have also suggested an enhanced freshwater contribution from the northern margin of the EMS during this period, attributed to high precipitation driven by westerly winds and the melting of glacial ice-sheets in response to Northern Hemisphere climate changes (e.g., Emeis et al., 2003; Filippidi et al., 2016; Kallel et al., 1997; Melki et al., 2009; Toucanne et al., 2015; Tesi et al., 2017). The aforementioned events would also have led to a reduction of deep-water formation in the EMS, thereby limiting the oxygen supply to the deep water (Rohling, 1994). Furthermore, it has been shown that the Mediterranean thermohaline circulation is sensitive to rapid climatic changes in the northern hemisphere (e.g., Heinrich and Dansgaard - Oeschger events) (Bartov et al., 2003; Colin et al., 2021; Kallel et al., 1997; Martrat et al., 2004; Rohling et al., 1998). Given this sensitivity of the Mediterranean thermohaline circulation to climate change at high and low latitudes, further research at a high temporal resolution is needed to understand how circulation dynamics may have evolved or fluctuated, particularly during the well-developed S5 deposition.

In this study, we have performed analyses of major and trace element concentrations based on XRF core scanning, δ¹⁸O of planktonic foraminifera and Nd isotope compositions (εNd) of mixed planktonic foraminifera samples from a core collected in the southeastern Aegean Sea characterized by a notably thick layer of sapropel S5 (150 cm laminated) and a core located in the Sardinia Channel. The aim was to better constrain the water mass circulation patterns of the last climatic cycle, focusing particularly on the Last Interglacial Maximum (MIS5e). Given a residence time of 500-1000 years (Siddall et al., 2008; Tachikawa et al., 2003) and distinct local
basin-scale sources, εNd can be considered to behave quasi-conservatively, thereby offering significant potential for fingerprinting the provenance of Mediterranean water masses at the present time (Garcia-Solsona et al., 2020; Montagna et al., 2022) and in the past (Cornuault et al., 2018; Colin et al., 2021; Dubois-Dauphin et al., 2017; Duhamel et al., 2020; Wu et al., 2019). Modern Mediterranean seawater displays a wide range of εNd values, from -11 to -4.2, with unradiogenic surface water entering the Mediterranean Sea as Atlantic Water and more radiogenic intermediate and deep waters originating from the Levantine Basin in contact with radiogenic sediments from the Nile River and the Egyptian/Libyan margin (Garcia-Solsona et al., 2020; Henry et al., 1994; Montagna et al., 2022; Tachikawa et al., 2004; Vance et al., 2004). The εNd record from the southeastern Aegean Sea, presented in this study, when compared with prior results from the EMS (Duhamel et al., 2020; Osborne et al., 2008; Scrivner et al., 2004) allows us to constrain the hydrological circulation during sapropel S5 at high temporal resolution and to investigate the link with climate changes at high and low latitudes.

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## 2. Hydrological setting

The semi-enclosed Mediterranean Sea is located between the subtropical high-pressure belt over North Africa and the temperate westerlies over central and southern Europe. Given its physiography, the Mediterranean Sea has a thermohaline circulation that is highly sensitive to climate changes in both high and low latitudes (e.g., Colin et al., 2021; Emeis et al., 2003; Kallel et al., 1997; Melki et al., 2009; Pinardi and Masetti, 2000; Roether et al., 1996; Rohling et al., 2002, 2015; Toucanne et al., 2015; Tesi et al., 2017;). The combination of low-latitude climate-related excess evaporation and cold spells originating from mid-latitudes results in the frequent formation of deep-water (Lolis et al., 2002). This phenomenon is observed in the two northernmost marginal seas of the EMS, i.e. the Adriatic Sea and the Aegean Sea (Casford et al., 2003; Maheras et al., 1999; Poulos et al., 1997; Saaroni et al., 1996).

In details, the Mediterranean Sea exchanges surface and intermediate waters with the Atlantic Ocean through the Gibraltar Strait (sill depth ~ 300 m) and surface water with the Black Sea through the Dardanelles Strait (sill depth ~100 m) (Bryden and Stommel, 1984; Béthoux, 1989; Béranger et al., 2005, 2010; Milot et al., 1995, 2005; Pistek et al., 1985; Roether and Well, 2001; ).
A negative freshwater budget drives the formation of an anti-estuarine circulation in the Mediterranean Sea (POEM Group, 1992; Pinardi and Masetti, 2000; Roether et al., 1996;). Hence, the relatively fresh surface Atlantic Water (AW) (salinity ~ 36.5), which has an unradiogenic εNd of ~9.7 (Tachikawa et al., 2004), flows into the WMS and is then modified by mixing with the ambient surface water, reaching a value of ~10.4 in the Alboran Sea (Spivack and Wasserburg, 1988; Tachikawa et al., 2004). The Modified Atlantic Water (MAW) flows eastward at 50-200 m water depth with εNd values from -10.8 to -9.0 in the WMS (Henry et al., 1994) and from -9.8 to -4.9 in the EMS (Tachikawa et al., 2004; Vance et al., 2004). Given the prevailing conditions of evaporation exceeding precipitation and river runoff, the MAW becomes progressively saltier (~38.5-37.5) as it circulates in the EMS.

During wintertime, the combination of intense cooling and strong wind-induced heat loss results in the formation of denser waters that sink via convection, creating intermediate waters in the Aegean Sea and the Levantine Basin as well as deep waters in the Gulf of Lions and the Adriatic Sea (Robinson et al., 2001; Schroeder et al., 2016). In particular, the Levantine Intermediate Water (LIW) is formed in the Cyprus-Rhodes area and it spreads westwards throughout the entire Mediterranean Basin at ~150-700 m water depth (Fig. 1) (Lascaratos et al., 1993; Malanotte-Rizzoli et al., 1999). The LIW acquires its Nd isotope composition primarily from the partial dissolution of particles originating from the Nile River and the Egyptian/Libyan margin sediments of the EMS, which exhibit radiogenic εNd values ranging from -7 to -2 (Ayache et al., 2016; Montagna et al., 2022; Tachikawa et al., 2004; Weldeab et al., 2002). This more radiogenic εNd signature is explained by sediment originating from the erosion of Ethiopian traps, including Cenozoic basaltic silicate rocks (εNd > 0) (e.g., Garzanti et al., 2015). The LIW is thus characterized by a radiogenic εNd of ~4.8 in the eastern part of the EMS and by an east-west εNd gradient ranging from -4.8 to -9.2, resulting from mixing with more unradiogenic overlying and underlying water masses along its path (Henry et al., 1994; Montagna et al., 2022; Tachikawa et al., 2004; Vance et al., 2004).

In the Adriatic Sea, LIW participates in the formation of Adriatic Deep Water (AdDW), which flows into the deep Ionian Basin and contributes, along with Aegean Deep Water (AeDW), to the formation of the Eastern Mediterranean Deep Water (EMDW) (e.g., Milner et al., 2012)
(Fig. 1). The εNd values of AdDW and AeDW are -7.40±0.15 and -6.10±0.16, respectively (Montagna et al., 2022). The formation of the Western Mediterranean Deep Water (WMDW) occurs in the Gulf of Lions during the winter months, resulting from the mixing of relatively fresh surface water with the saline LIW; it then spreads into the Balearic Basin and Tyrrhenian Sea between ~ 2000 and ~3000 m (Millot, 1999; Schroeder et al., 2016) (Fig. 1). The WMDW is characterized by an average εNd value of -8.60±0.32 (n=4; 2σ) (Garcia-Solsona et al., 2020; Montagna et al., 2022). Between the WMDW and the LIW (from ~700 to ~ 2000 m), the Tyrrhenian Deep Water (TDW) (Millot et al., 2005) has an average εNd value of ~-9.01±0.50 (n=3 2σ) (Garcia-Solsona et al., 2020; Montagna et al., 2022).

<figure><img src="image_2.png" /><figcaption>**Fig. 1** Bathymetric map of the Mediterranean Sea showing locations of cores MD90-920 (35.58°N, 28.46 E, water depth 1620 m) and MD90-912 (38°18.67′N, 9°46.12′E, water depth 1020 m) (red-dots) and all other sites discussed in this study (black dots). The simplified circulation pattern of the Levantine Intermediate Water (grey arrows) and deep water (dashed black arrows) is modified from Pinarti and Masetti (2000) and Schroeder et al. (2012). EMDW: Eastern Mediterranean Deep Water; AeDW: Aegean Deep Water; AdDW: Adriatic Deep Water; WMDW: Western Mediterranean Deep Water; LIW: Levantine Intermediate Water.</figcaption></figure>

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## 3. Material and methods

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### 3.1 Material
Cores MD90-920 (35.58°N, 28.46°E, water depth 1620 m) and MD90-912 (38.31°N, 9.77°E, water depth 1020 m) were collected in 1990 in the southeastern Aegean Sea and the Tyrrhenian Sea (Sardinia Channel), respectively, during the PROMETE Legs I and III campaign aboard the R/V *Marion Dufresne* (Fig. 1). Core MD90-920 is located on the pathway of the outflow of the AeDW from the Aegean Sea to the Levantine Basin. The upper 1000 cm of core MD90-920 consists of pale cream to yellowish brown foraminiferal and nannofossil ooze, interbedded with 4 sapropel layers (S1, S3, S4 and S5). Core MD90-912 is located in the upper portion of TDW, which results from the mixing of the WMDW, LIW and upper EMDW (Fig. 1). The core consists of a homogeneous brownish, muddy, carbonate-rich sediment.

## 3.2 Methods

### 3.2.1. Planktonic foraminifera $\delta^{18}$O analyses

$\delta^{18}$O analyses were performed on the planktonic foraminifera *Globigerinoides ruber* in 77 samples from core MD90-920 and on the planktonic foraminifera *Globigerinoides bulloides* in 45 samples from core MD90-912 in order to cover one climatic cycle. From each sample, approximatively 4-8 clean and well-preserved shells (250-315 µm) were hand-picked. For core MD90-920, $\delta^{18}$O analyses were carried out on a Finnigan MAT-251 mass spectrometer at the Laboratoire des Sciences du Climat et de l'Environnement (LSCE, France). For core MD90-912, samples were analyzed at the State Key Laboratory of Marine Geology (Tongji University) using a Finnigan MAT-253 mass spectrometer. The oxygen isotope compositions are expressed in conventional delta notation ($\delta$) relative to the isotopic ratio of the carbon dioxide gas derived from the Vienna Pee Dee Belemnite (VPDB) standard. Ratios are reported in ‰. The mean external reproducibility of carbonate standards is better than $\pm$0.05‰ for $\delta^{18}$O for both mass spectrometers used in this study.

### 3.2.2 XRF core scanning

The bulk XRF geochemical data were obtained on core MD90-920 directly on the surfaces of split sections at 1 cm resolution using an *Avaatech™* XRF core scanner at the Laboratoire des Sciences du Climat et de l'Environnement (France). To prevent contamination of the XRF measurement unit and desiccation of the sediment, the core surface was covered with a 4 µm thin Ultralene foil. To ensure the best quality data, the sediment surface was carefully smoothed, with
all air bubbles under the foil and all wrinkles in the foil carefully eliminated. The measurement
was performed under conditions of 10 KeV, 30 KeV and 50 KeV to obtain intensities of elements
from Al to Fe. In this study, only elements with an average total count exceeding 1000 were
considered.

### 3.2.3 Foraminiferal Nd isotope compositions

Neodymium isotope compositions were measured on ~30 mg of mixed planktonic
foraminifera from the >150 µm size fraction, without employing any oxidative-reductive cleaning
procedure, as this approach has been demonstrated to be suitable for extracting deep-water Nd
isotope compositions (e.g., Tachikawa et al., 2014, Wu et al., 2019). The foraminifera samples
were gently crushed between two glass slides under the microscope to open all chambers, and
then ultrasonicated for 1 min in MilliQ water before the suspended particles were pipetted off.
This step was repeated until the water was clear and free of clay particles. The remaining
foraminifera were dissolved using weak leaching with stepwise 100 µl nitric acid (0.5 M HNO₃).
The dissolved samples were centrifuged, and the supernatant was immediately transferred to
Teflon beakers to prevent leaching of any possible remaining solid phases.

Neodymium was then separated following the analytical procedure of Copard et al. (2010).
In brief, samples were loaded in 2 ml of 1 M HNO₃ into pre-conditioned columns (with 83 mg of
TRU spec™ resin). Unwanted cations were eluted using five portions of 0.5 ml of 1 M HNO₃. The
TRU spec™ columns were then placed over the LN spec™ columns. The light REEs were eluted
from the upper columns to the LN spec™ columns using seven portions of 0.05 M HNO₃. After
decoupling from the TRU spec™ columns, La, Ce, and most of the Pr were rinsed from the LN
spec™ columns with 2.5 ml of 0.25 M HCl. Nd was then eluted with an additional 3.25 ml of 0.25
M HCl. The cleaning procedure and purification of Nd were carried out in a class 100 clean
laboratory using ultrapure reagents.

The ¹⁴³Nd/¹⁴⁴Nd ratios of purified Nd fractions were measured by Multi-Collector
Inductively Coupled Plasma Mass Spectrometry (MC-ICP-MS, Thermo Fisher NeptunePlus) at
PANOPLY's analytical facilities at the Laboratoire des Sciences du Climat et de l'Environnement,
University Paris-Saclay, France. The solutions were analyzed at a concentration of 10 to 15 ppb.
The instrumental mass fractionation was corrected by normalizing ¹⁴⁶Nd/¹⁴⁴Nd to 0.7219
(O’Nions et al., 1977) and applying an exponential fractionation law. During the analysis, every group of three samples was bracketed with the JNdi-1 standard with Nd concentrations similar to those of the samples. Replicate analyses of the JNdi-1 standards yielded mean ¹⁴³Nd/¹⁴⁴Nd ratios of 0.512108 ± 0.000009 (2σ, n=32), closely matching the accepted value of 0.512115 ± 0.000006 (Tanaka et al., 2000). The external reproducibility (2σ) of εNd values deduced from repeated measurements of the JNdi-1 standard was around 0.2 and 0.4 εNd units for the different analytical sessions. The reported analytical uncertainty is the combination of the external reproducibility of the within-session standards and the internal measurement error of each sample. Total blanks were <50 pg and can be ignored as they represent <0.1% of the sample characterized by the lowest amount of Nd. Neodymium isotopic compositions are expressed as εNd = [(¹⁴³Nd/¹⁴⁴Nd)sample/(^143Nd/^144Nd)CHUR-1]×10000, with the present-day (¹⁴³Nd/¹⁴⁴Nd)CHUR of 0.512638 (Jacobsen and Wasserburg 1980).

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## 4. Results

### 4.1 Chronological framework

The chronologies of the studied cores were based on the δ¹⁸O stratigraphy of the planktonic foraminifera *G. ruber* for core MD90-920 and *G. bulloides* for core MD90-912. The δ¹⁸O records of planktonic foraminifera were correlated with that of the ODP site 967 (Grant et al., 2022) using the *Analyseries* software (Paillard et al., 1996) (Fig. 2).

The chronology of the upper 125 cm of core MD90-912 was previously published and established through linear interpolation from four calibrated AMS ¹⁴C dates obtained from planktonic foraminifer *G. bulloides* (Colin et al., 2021) (Fig. 2c). The new δ¹⁸O values obtained on *G. bulloides* have permitted the refinement of the age model for core MD90-912 for the last climatic cycle (Fig. 2c). The correlation of the *G. bulloides* δ¹⁸O record to that of the LR04 record indicates that the upper 3.13 m of core MD90-912 cover the last climatic cycle with a mean sedimentation rate of 2.08 cm/ka (Fig. 2c). The δ¹⁸O values of *G. bulloides* ranged between 1.1 and 3.9 ‰, with the highest and lowest values recorded during the MIS2 and the MIS5, respectively.

The *G. ruber* δ¹⁸O record for core MD90-920 displays a glacial-interglacial difference (based
on maximum and minimum values) of ~5.5 ‰ (Fig. 2b). The age model indicates that the upper
10.30 m of the core encompass the last 145 ka, with a linear mean sedimentation rate of 7.1 cm/ka.
The core displays dark layers, characterized by low L* values (Fig. 3a), enabling the identification
of sapropels S1, S3, S4 and S5 (Fig. 3). Sapropel S5 is the thickest of the four, measuring 150 cm,
whereas S1, S3 and S4 reach 15, 10 and 10 cm, respectively. The δ¹⁸O of *G. ruber* gradually
increased from -1.68 ‰ during the last interglacial period to 3.91 ‰ during the last glacial
maximum (Fig. 3b). The δ¹⁸O values throughout the sapropels are generally low (Fig. 3b).

<figure><img src="image_3.png" /><figcaption>**Fig. 2** Age models of cores MD90-920 and MD90-912 based on planktonic foraminifer oxygen isotopes (δ¹⁸O) and radiocarbon dates. (a) δ¹⁸O of *G. ruber* from core ODP site 967 (Grant et al., 2022); (b) δ¹⁸O of *G. ruber* from core MD90-920; (c) δ¹⁸O of *G. bulloides* from core MD90-912;</figcaption></figure>
triangles correspond to ¹⁴C dating. The boundaries of the Marine Isotope Stages (MIS) in this figure, as well as in other figures throughout this study, have been defined according to Lisiecki and Raymo (2005).

<figure><img src="image_4.png" /></figure>
**Fig.3** (a) δ¹⁸O curve obtained on planktonic foraminifera G. ruber of core MD90-920 compared with color reflectance parameters L* (higher L* values indicate lighter colors); (b) εNd records obtained on mixed planktonic foraminifera from core MD90-920; (c) XRF Zr/Rb ratio of core MD90-920; (d) XRF Ba/Al ratio of core MD90-920; (e) XRF S/Al ratio of core MD90-920. The grey bars represent the position of sapropels S1, S3, S4 and S5 in core MD90-920. Due to post-depositional reoxidation processes that often follow anoxic conditions, identifying and accurately measuring the thickness of sapropels in sedimentary sequences is challenging and the color of the upper portion of the sapropel from black to grey thereby obscuring its visual identification. Consequently, the thicknesses of sapropels, indicated by the grey bars have been determined based on variations in Ba/Al and S/Al following the method used by Filippidi et al., 2019 and presented in detail in the supplementary text.

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## 4.2 XRF core scanner

Sapropels S3, S4 and S5 consistently exhibited higher Ba/Al and S/Al ratios (Fig. 3d and 3e). In addition, sapropels also coincide with lower Zr/Rb values (Fig. 3c). Notably, the well-developed sapropel S5 allowed us to examine in detail geochemical variations during one of the most significant organic carbon preservation events of the late Quaternary (Fig. 3). At the onset of sapropel S5 (~128 ka), the increase in Ba/Al and S/Al ratios is associated with a decrease in δ¹⁸O of G. ruber and in Zr/Rb ratio (Fig. 4). The first part of the S5 (128-123 ka) is associated to steady low values for Zr/Rb, and high values for Ba/Al as well as a slight increase trend for S/Al ratio. During the second part of the S5 (123-120 ka), Ba/Al and S/Al ratios exhibited a slight increase, whereas the Zr/Rb ratio displayed a marked continuous increase. Finally, the δ¹⁸O of G. ruber remains low and relatively stable during the sapropel S5 (Fig. 4).

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<figure><img src="image_5.png" /><figcaption>**Fig. 4** Sapropel S5 of core MD90-920: (a) δ¹⁸O curve obtained on G. ruber; (b) foraminiferal εNd record; (c) XRF Zr/Rb ratio; (d) XRF Ba/Al ratio; (e) XRF S/Al ratio. The vertical orange bar is the early stage of S5 and the blue bar is the late stage of S5.</figcaption></figure>
## 4.3 Foraminiferal εNd records

The foraminiferal εNd record of core MD90-920 displays a significant range of values, from -5.75±0.16 to -2.41±0.15, during the last climatic cycle (Table 1 and Fig. 3b). Overall, long-term variations of εNd values did not show clear glacial-interglacial fluctuations, but rather exhibited variability within the deposition of sapropels S3, S4 and S5 (Fig. 2b). Each sapropel deposit was associated with a slight increase in the εNd values, observed at ~120 ka, ~109 ka and ~82-78 ka. At the early stage of S5 (~129-127 ka), εNd values decreased from -3.70±0.28 to -5.08±0.35 (Fig. 4b), whereas during the later phase of S5 (around ~120 ka), εNd values exhibited the most radiogenic levels, reaching -2.41±0.15 (Fig. 4b).

For core MD90-912, εNd values ranged from -6.83±0.35 to -8.88±0.18 during the last climatic cycle (Table 1 and Fig. 5). In general, each warm substage of the MIS5, which is associated with sapropel S3, S4 and S5 deposition in the EMS, is characterized by an increase in εNd values, with the greatest increases occurring at ~121 ka, ~102 ka and ~81 ka (Fig. 5b).

<figure><img src="image_6.png" /><figcaption>**Fig. 5** (a) $\delta^{18}$O curve obtained on planktonic foraminifera G. bulloides of core MD90-912; (b) εNd records obtained on mixed planktonic foraminifera from core MD90-912. The dash line indicates the time interval of sapropel deposition in the EMS.</figcaption></figure>
<table><thead><tr><th>Depth (cm)</th><th>Age (ka)</th><th><sup>143</sup>Nd/<sup>144</sup>Nd</th><th>± 2σ</th><th>εNd</th><th>± 2σ</th></tr></thead><tbody><tr><td colspan="6"><strong>MD90-920 (35.58° N, 28.46 E, 1620 m)</strong></td></tr><tr><td>70</td><td>8.38</td><td>0.512405</td><td>0.000008</td><td>-4.55</td><td>0.15</td></tr><tr><td>90</td><td>9.54</td><td>0.512409</td><td>0.000008</td><td>-4.47</td><td>0.16</td></tr><tr><td>100</td><td>9.89</td><td>0.512403</td><td>0.000009</td><td>-4.58</td><td>0.17</td></tr><tr><td>120</td><td>10.58</td><td>0.512419</td><td>0.000024</td><td>-4.27</td><td>0.47</td></tr><tr><td>130</td><td>10.93</td><td>0.512391</td><td>0.000042</td><td>-4.82</td><td>0.82</td></tr><tr><td>185</td><td>16.23</td><td>0.512397</td><td>0.000011</td><td>-4.71</td><td>0.22</td></tr><tr><td>250</td><td>23.38</td><td>0.512441</td><td>0.000011</td><td>-3.85</td><td>0.21</td></tr><tr><td>340</td><td>33.29</td><td>0.512374</td><td>0.000028</td><td>-5.15</td><td>0.54</td></tr><tr><td>350</td><td>34.39</td><td>0.512366</td><td>0.000015</td><td>-5.30</td><td>0.29</td></tr><tr><td>390</td><td>38.79</td><td>0.512380</td><td>0.000008</td><td>-5.03</td><td>0.16</td></tr><tr><td>440</td><td>45.86</td><td>0.512420</td><td>0.000010</td><td>-4.25</td><td>0.19</td></tr><tr><td>450</td><td>47.49</td><td>0.512380</td><td>0.000029</td><td>-5.03</td><td>0.57</td></tr><tr><td>480</td><td>52.36</td><td>0.512437</td><td>0.000011</td><td>-3.92</td><td>0.22</td></tr><tr><td>520</td><td>58.85</td><td>0.512402</td><td>0.000011</td><td>-4.61</td><td>0.22</td></tr><tr><td>610</td><td>74.41</td><td>0.512366</td><td>0.000011</td><td>-5.30</td><td>0.22</td></tr><tr><td>640</td><td>82.15</td><td>0.512432</td><td>0.000027</td><td>-4.01</td><td>0.52</td></tr><tr><td>650</td><td>84.73</td><td>0.512409</td><td>0.000012</td><td>-4.46</td><td>0.24</td></tr><tr><td>660</td><td>88.11</td><td>0.512393</td><td>0.000008</td><td>-4.78</td><td>0.16</td></tr><tr><td>700</td><td>101.64</td><td>0.512444</td><td>0.000027</td><td>-3.79</td><td>0.54</td></tr><tr><td>720</td><td>108.41</td><td>0.512391</td><td>0.000016</td><td>-4.81</td><td>0.31</td></tr><tr><td>730</td><td>111.79</td><td>0.512352</td><td>0.000007</td><td>-5.58</td><td>0.13</td></tr><tr><td>740</td><td>115.18</td><td>0.512343</td><td>0.000008</td><td>-5.75</td><td>0.16</td></tr><tr><td>770</td><td>120.53</td><td>0.512515</td><td>0.000007</td><td>-2.41</td><td>0.15</td></tr><tr><td>780</td><td>121.06</td><td>0.512445</td><td>0.000009</td><td>-3.76</td><td>0.17</td></tr><tr><td>800</td><td>122.12</td><td>0.512396</td><td>0.000009</td><td>-4.73</td><td>0.17</td></tr><tr><td>810</td><td>122.65</td><td>0.512381</td><td>0.000029</td><td>-5.02</td><td>0.56</td></tr><tr><td>820</td><td>123.18</td><td>0.512404</td><td>0.000009</td><td>-4.57</td><td>0.17</td></tr><tr><td>830</td><td>123.71</td><td>0.512364</td><td>0.000038</td><td>-5.35</td><td>0.74</td></tr><tr><td>850</td><td>124.76</td><td>0.512393</td><td>0.000009</td><td>-4.78</td><td>0.18</td></tr><tr><td>870</td><td>125.82</td><td>0.512424</td><td>0.000008</td><td>-4.18</td><td>0.15</td></tr><tr><td>880</td><td>126.35</td><td>0.512414</td><td>0.000008</td><td>-4.38</td><td>0.15</td></tr><tr><td>890</td><td>126.88</td><td>0.512398</td><td>0.000008</td><td>-4.67</td><td>0.15</td></tr><tr><td>900</td><td>127.41</td><td>0.512377</td><td>0.000018</td><td>-5.08</td><td>0.35</td></tr><tr><td>910</td><td>127.94</td><td>0.512394</td><td>0.000024</td><td>-4.76</td><td>0.46</td></tr><tr><td>920</td><td>128.47</td><td>0.512428</td><td>0.000008</td><td>-4.09</td><td>0.16</td></tr><tr><td>930</td><td>128.86</td><td>0.512448</td><td>0.000014</td><td>-3.70</td><td>0.28</td></tr><tr><td>970</td><td>131.97</td><td>0.512431</td><td>0.000011</td><td>-4.04</td><td>0.22</td></tr><tr><td>1120</td><td>144.61</td><td>0.512418</td><td>0.000021</td><td>-4.28</td><td>0.42</td></tr><tr><td colspan="6"><strong>MD90-912 (38°18.67′N, 9°46.12′E, 1020 m)</strong></td></tr><tr><td>2</td><td>0.43</td><td>0.512209</td><td>0.000015</td><td>-8.37</td><td>0.29</td></tr></tbody></table>
<table><caption>**Table 1:** Nd isotope composition obtained on mixed planktonic foraminifera samples from cores</caption><tr><td>9</td><td>4.23</td><td>0.512208</td><td>0.000022</td><td>-8.39</td><td>0.43</td></tr><tr><td>15</td><td>8.13</td><td>0.512223</td><td>0.000016</td><td>-8.10</td><td>0.31</td></tr><tr><td>19</td><td>10.73</td><td>0.512244</td><td>0.000013</td><td>-7.69</td><td>0.25</td></tr><tr><td>35</td><td>16.68</td><td>0.512192</td><td>0.000011</td><td>-8.70</td><td>0.21</td></tr><tr><td>45</td><td>20.14</td><td>0.512197</td><td>0.000017</td><td>-8.60</td><td>0.33</td></tr><tr><td>55</td><td>24.04</td><td>0.512209</td><td>0.000008</td><td>-8.37</td><td>0.16</td></tr><tr><td>69</td><td>29.50</td><td>0.512219</td><td>0.000008</td><td>-8.17</td><td>0.16</td></tr><tr><td>79</td><td>33.39</td><td>0.512218</td><td>0.000011</td><td>-8.19</td><td>0.21</td></tr><tr><td>89</td><td>37.29</td><td>0.512229</td><td>0.000010</td><td>-7.98</td><td>0.20</td></tr><tr><td>99</td><td>41.19</td><td>0.512230</td><td>0.000011</td><td>-7.96</td><td>0.21</td></tr><tr><td>115</td><td>47.42</td><td>0.512218</td><td>0.000011</td><td>-8.19</td><td>0.21</td></tr><tr><td>125</td><td>53.11</td><td>0.512207</td><td>0.000009</td><td>-8.41</td><td>0.18</td></tr><tr><td>135</td><td>57.19</td><td>0.512222</td><td>0.000009</td><td>-8.11</td><td>0.18</td></tr><tr><td>145</td><td>60.64</td><td>0.512223</td><td>0.000009</td><td>-8.10</td><td>0.18</td></tr><tr><td>155</td><td>64.10</td><td>0.512227</td><td>0.000007</td><td>-8.02</td><td>0.14</td></tr><tr><td>165</td><td>74.61</td><td>0.512213</td><td>0.000009</td><td>-8.29</td><td>0.18</td></tr><tr><td>175</td><td>78.58</td><td>0.512181</td><td>0.000009</td><td>-8.91</td><td>0.18</td></tr><tr><td>185</td><td>82.55</td><td>0.512231</td><td>0.000013</td><td>-7.94</td><td>0.25</td></tr><tr><td>190</td><td>84.53</td><td>0.512212</td><td>0.000009</td><td>-8.31</td><td>0.18</td></tr><tr><td>200</td><td>90.12</td><td>0.512195</td><td>0.000009</td><td>-8.64</td><td>0.18</td></tr><tr><td>205</td><td>92.96</td><td>0.512230</td><td>0.000010</td><td>-7.96</td><td>0.20</td></tr><tr><td>210</td><td>95.80</td><td>0.512210</td><td>0.000009</td><td>-8.35</td><td>0.18</td></tr><tr><td>215</td><td>98.63</td><td>0.512205</td><td>0.000014</td><td>-8.45</td><td>0.27</td></tr><tr><td>220</td><td>101.17</td><td>0.512218</td><td>0.000009</td><td>-8.19</td><td>0.18</td></tr><tr><td>225</td><td>102.99</td><td>0.512259</td><td>0.000013</td><td>-7.39</td><td>0.25</td></tr><tr><td>230</td><td>104.80</td><td>0.512245</td><td>0.000010</td><td>-7.67</td><td>0.20</td></tr><tr><td>235</td><td>106.62</td><td>0.512232</td><td>0.000007</td><td>-7.92</td><td>0.14</td></tr><tr><td>240</td><td>114.60</td><td>0.512223</td><td>0.000017</td><td>-8.10</td><td>0.33</td></tr><tr><td>245</td><td>116.43</td><td>0.512225</td><td>0.000009</td><td>-8.06</td><td>0.18</td></tr><tr><td>255</td><td>121.57</td><td>0.512288</td><td>0.000018</td><td>-6.83</td><td>0.35</td></tr><tr><td>260</td><td>123.33</td><td>0.512202</td><td>0.000009</td><td>-8.51</td><td>0.18</td></tr><tr><td>265</td><td>123.98</td><td>0.512225</td><td>0.000010</td><td>-8.06</td><td>0.20</td></tr><tr><td>270</td><td>124.63</td><td>0.512206</td><td>0.000009</td><td>-8.43</td><td>0.18</td></tr><tr><td>275</td><td>125.28</td><td>0.512239</td><td>0.000013</td><td>-7.78</td><td>0.25</td></tr><tr><td>280</td><td>125.93</td><td>0.512219</td><td>0.000010</td><td>-8.17</td><td>0.20</td></tr><tr><td>285</td><td>127.07</td><td>0.512200</td><td>0.000010</td><td>-8.54</td><td>0.20</td></tr><tr><td>290</td><td>128.26</td><td>0.512211</td><td>0.000011</td><td>-8.33</td><td>0.21</td></tr><tr><td>295</td><td>129.45</td><td>0.512200</td><td>0.000009</td><td>-8.54</td><td>0.18</td></tr><tr><td>305</td><td>133.25</td><td>0.512187</td><td>0.000009</td><td>-8.80</td><td>0.18</td></tr><tr><td>311</td><td>137.11</td><td>0.512183</td><td>0.000009</td><td>-8.88</td><td>0.18</td></tr><tr><td>315</td><td>139.17</td><td>0.512188</td><td>0.000010</td><td>-8.78</td><td>0.20</td></tr></table>
340 *MD90-912 and MD90-920. The uncertainties are given at the two-sigma (2σ) level for the internal*
341 *and external error.*

342

343 **5. Discussion**

344 **5.1 Significance of the authigenic εNd variations during the last climatic cycle**

345 Several studies have shown that the neodymium isotope composition of uncleaned planktonic
346 foraminifera reflects the dissolved εNd composition of bottom water and/or ambient pore water
347 due to post-depositional incorporation of Nd into authigenic Fe-Mn oxyhydroxide coatings (e.g.,
348 Tachikawa et al., 2014; Wu et al., 2019). Montagna et al. (2022) recently compiled an extensive
349 εNd dataset for the Mediterranean Sea, revealing a clear εNd-salinity correlation that indicates
350 that dissolved εNd behaves overall conservatively in the Mediterranean Sea and serves as a
351 reliable tracer for differentiating among various water masses. Unfortunately, the foraminiferal
352 εNd of the core-top sample from core MD90-920 is not available, which precludes a direct
353 comparison with modern seawater εNd. Nevertheless, the εNd value of the core-top sample of
354 core LC21 (-5.05±0.33; Osborne et al., 2010), collected at 1522 m water depth in proximity to
355 core MD90-920, is similar to the modern seawater value at station 64PE370-28 (Fig. 1) (εNd
356 value at 1000 m of -5.30±0.46; Montagna et al., 2022). For core MD90-912, the foraminiferal
357 εNd of the core-top sample is -8.37±0.3 (Table 1 and Fig. 6e), which is similar to that of the
358 adjacent station Meteor-317 (Fig. 1) (εNd value at 3268 m of -8.70±0.50; Montagna et al., 2022).
359 This indicates that foraminifera εNd mainly records the signature of bottom water for the two
360 studied sites. However, Montagna et al. (2022) also pointed out that the εNd of the water masses
361 can be modified through the input of lithogenic Nd from dust and river particles to the
362 Mediterranean Sea, as well as interaction with radiogenic and unradiogenic sediments (boundary
363 exchange). Saharan dusts are characterized by unradiogenic εNd values (-15.4 to -11; Scheuvens
364 et al., 2013) and may have contributed to the unradiogenic Nd signature observed in surface and
365 intermediate water masses of the Mediterranean Sea (e.g. Arsouze et al., 2009; Ayache et al., 2023;
366 Rempfer et al., 2011). The present-day distribution of εNd in the surface and intermediate water
367 masses of the Mediterranean Sea is mainly attributed to the mixing of the AW and the LIW,
368 suggesting a negligible contribution of lithogenic Nd input from eolian dust (Montagna et al.,
This can be mainly attributed to the relatively low Nd concentration and fractional solubility of eolian dust (Tachikawa et al., 2004).

<figure><img src="image_7.png" /></figure>
**Fig. 6** (a) Comparison between δ¹⁸O curve obtained on G. ruber (blue curve) of core MD90-920 and June insolation received by the Earth at 15°N (Laskar et al., 2004); (b) δ¹⁸O values of speleothem of Soreq Cave in Israel (blue line, Bar-Matthews et al., 2003) and the wet/dry proxy of northern Levantine Basin (purple line, ODP site 967, Grant et al., 2017); εNd records obtained on mixed planktonic foraminifera from (c) core MD90-920 (this study); (d) core MD90-964 (Duhamel et al., 2020); (e) core MD90-912 (this study); (f) core MD21-2472 (Colin et al., 2021); (g) variations in the Relative Sea Level (RSL) (light-blue line, Grant et al., 2012). The shadow gray bars represent the position of the sapropels, and the dashed line for the core MD90-912 and MD01-2472 is an equivalent time of sapropel deposits. The pink horizontal bars indicate the modern seawater εNd values, measured close to the core sites (Tachikawa et al., 2004; Garcia-Solsona et al., 2020). NBEM: the Northern Borderlands of the Eastern Mediterranean.

However, glacial periods were accompanied by an increase in aridity across North Africa and a two-fold increase in the influx of Saharan dust into the Mediterranean Sea (e.g., Crocker et al., 2022; Ehrmann et al., 2017; Skonieczny et al., 2019). Such increases in dust inputs do not appear to have significantly modified the seawater εNd values reconstructed from both cores MD90-920 and MD90-912 since glacial MIS are not associated with more unradiogenic foraminiferal εNd (Figs. 6c and 6e). Similarly, foraminiferal εNd records obtained in the Corsica Trough (core MD01-2472) and in the southern Levantine Basin (core MD90-964) do not show more unradiogenic εNd values during the glacial MIS compared to interglacials MIS. In fact, the values are even more radiogenic during the glacial MIS for core MD90-964 (Colin et al., 2021; Duhamel et al., 2020) (Fig. 6). These findings are in agreement with other previous εNd records obtained from the WMS (Colin et al., 2021; Dubois-Dauphin et al., 2017; Jiménez-Espejo et al., 2015) and EMS (Cornuault et al., 2018; Wu et al., 2019), including those from cores located at intermediate water depth, suggesting that the Saharan dust had little impact on the past seawater εNd record of deep-water masses in the Mediterranean Sea.

Conversely, the distribution of εNd in the oceans can be significantly influenced by lithogenic Nd input from rivers (e.g., Rousseau et al., 2015; Yu et al., 2017). The Black Sea and the rivers along the northern borderlands of the EMS have been identified as the primary detrital
and freshwater sources to the Aegean Sea (Grant et al., 2012; Lane-Serff et al., 1997; Rohling et al., 2015). The εNd value of the northern rivers adjacent to the Aegean Sea is around -9.0 (Weldeab et al., 2002), which does not adequately explain the more radiogenic values observed in core MD90-920, particularly during the deposition of sapropels (Fig. 3b). The partial dissolution of highly radiogenic Nile-derived volcanic material (εNd from -3 to +3, Tachikawa et al., 2004) is the primary source of radiogenic Nd in the modern Mediterranean basin (e.g., Freydier et al., 2001; Scrivner et al., 2004; Montagna et al., 2022). This, together with the radiogenic Nd signal carried by the paleo-rivers of northern North Africa in the past (εNd from -2 to -1, Osborne et al., 2008), has been considered an important contribution of radiogenic Nd in the Levantine Basin also during past humid episodes (e.g. S5; Osborne et al., 2008; 2010). Consequently, the LIW and MAW acquire a radiogenic signature from the Nile River. Such radiogenic εNd water masses are then transported westward across the EMS through the thermohaline circulation, contributing to the higher radiogenic values observed in cores MD90-920 and MD90-912 (Fig. 6c and 6e).

Core MD90-920 exhibits the most radiogenic foraminiferal εNd values during sapropel depositions, reaching up to -2.4 during sapropel S5 (Fig. 3b). Similar variations have been observed in core MD90-964, located near the Nile River mouth (33°02.75′N, 32°38.57′E, water depth of 1375 m, Figs. 1 and 6d), where the sapropels of the last climatic cycle are likewise associated with more radiogenic εNd values (Duhamel et al., 2020). Given that radiogenic sediments and freshwater discharges of the Nile River represent the primary source of radiogenic Nd to the EMS, the strong covariances between the εNd record and precession minima-related insolation maxima observed in core MD90-920 (Fig. 6a and 6c) suggest that the freshwater balance of the EMS is significantly controlled by changes in the supply of monsoon-fueled freshwater from North Africa (Fig. 6b) (Amies et al., 2019; Lisiecki et al., 2005; Rossignol-Strick et al., 1982; Rossignol-Strick, 1983, 1999).

Radiogenic εNd values during sapropel depositions are also observed in core MD90-912 from the Sardinia Channel (Fig. 6e), which track the main pathways of the upper EMDW from the Tyrrhenian Sea to the Balearic Sea. Indeed, the core MD90-912, bathed by the TDW, shows more radiogenic εNd values (reaching ~-7.5) during sapropel deposition in the EMS (Fig. 6e). It
is noteworthy that the modern mean εNd value for the TDW close to core MD01-2472 is ~
9.01±0.50 (x=3; 2σ) (Garcia-Solsona et al., 2020; Montagna et al., 2022). This observation
suggests that the radiogenic εNd maxima observed during the sapropel events resulted from the
mixing of radiogenic LIW and the upper EMDW, which flows westward through the Siculo-
Tunisian Strait, with the TDW. This is in agreement with previous studies that hypothesized the
transfer of highly radiogenic εNd seawater from the EMS to the WMS through the Siculo-Tunisian
Strait during sapropel depositions (e.g., Colin et al., 2021).

## 5.2 Paleoenvironmental variations at high temporal resolution within S5

Sapropels often exhibit visible laminations on a sub-millimeter to millimeter-scale. This
feature is particularly evident in sapropels found in regions where sedimentary sequences present
high sediment accumulation rates, as observed in both marine sediment and terrestrial outcrops
(e.g., Corselli et al., 2002; Moller et al., 2012; Tesi et al., 2024). For core MD90-920, sapropel S5
is characterized by a well-preserved laminated interval associated with a high Si/Al ratio peak
(Fig. S1). Microscope observation of sapropel S5 samples revealed a high abundance of diatoms.
Scanning electron microscopy studies of S5 in core LC21 identified intervals with clear lamina
couplets that comprise laminae of rhizosolenid diatoms alternating with laminae of mixed
assemblages (Moller et al., 2012). These laminae are attributed to annual deposition or ‘varves’
based on correlation with the seasonal succession of diatom floras in the Mediterranean Sea
(Kemp et al., 1999, 2000; Pearce et al., 1998). It can therefore be hypothesized that the high Si/Al
ratio observed in core MD90-920 is due to the predominant contribution of diatoms to the organic
matter sedimentation.

The core MD90-920 is ideally located within the outflow of the Aegean Deep Water (AeDW),
to reconstruct past changes in AeDW formation by using the Zr/Rb ratio as a proxy of deep-sea
current velocity (e.g., McCave et al., 1995; Tesi et al., 2024; Wu et al., 2020). Indeed, Zr is
commonly found in heavy minerals and is associated with the coarse-grained fraction of
siliciclastic sediments due to its resistance to physical and chemical weathering (Fralick and
Kronberg, 1997). In contrast, Rb is prevalent in fine-grained particles, including K-rich clay
minerals and K-feldspars (Dypvik and Harris, 2001; Fralick and Kronberg, 1997). Consequently,
variations in the Zr/Rb ratio are mainly related to sediment texture, particularly the sorted
fraction of siliciclastic material, which has been extensively applied as a proxy for bottom current velocity (Alonso et al., 2023; Dypvik and Harris, 2001; Nichols et al., 2020; Toyos et al., 2020; Tesi et al., 2024). The decrease of the Zr/Rb ratio at the onset of sapropel S5 is associated with a decrease in the G. ruber δ¹⁸O and an increase in the Ba/Al and S/Al ratios (Fig. 4). This pattern implies that the increase in freshwater input at the beginning of the sapropel S5 deposition (reflected by lower δ¹⁸O, see also supplementary text) is coeval with a decline in AeDW formation (lower Zr/Rb ratio), potentially leading to its collapse at ~128 ka (Fig. 4). During the late stage of sapropel S5, the Zr/Rb ratio steadily increased, indicating a progressive intensification of the bottom current velocity, starting at ~123 ka (Fig. 4c), which is also associated with an abrupt change in εNd from -4.57±0.17 to -2.41±0.15. This suggests a major change in the origin of water masses in the northern Levantine Basin during the later stage of S5. Core LC21, located in the southern Aegean Sea nearby core MD90-920 (Fig. 1), displays a similar range and trend of εNd during sapropel S5 (Fig. 7a) (Osborne et al., 2010). The radiogenic peak in εNd observed at the end of sapropel S5 in core MD90-920 was also detected in core LC21, pointing to a regional εNd signal for the EMDW in the northern Levantine Sea, which suggests a major hydrological change discussed below with a compilation of past-seawater εNd records obtained previously during the S5 in the Levantine Basin.

475

476

## 5.3 The hydrological circulation pattern of the Levantine Basin during the deposition of S5

477

478

479

480

481

482

483

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εNd records from cores MD90-920, MD90-912, MD90-964, MD01-2472, LC21 and ODP site 967C are shown in figure 7 for a comparative analysis, with the aim of constraining the dynamics of deep-water masses circulation in the Levantine Basin during sapropel S5 (Colin et al., 2021; Duhamel et al., 2020; Osborne et al., 2008; Scrivner et al., 2004). It should be noted that all records have been transferred onto the ODP site 967 age model (Fig.S2) (Grant et al., 2022). Foraminiferal εNd records from core LC21 and ODP site 967C were obtained from planktonic foraminifera cleaned using the reductive-oxidative method outlined by Vance et al. (2004) and were interpreted as a surface water εNd signal (Osborne et al., 2008, 2010; Scrivner et al., 2004). However, recent studies have shown that the εNd values derived from both cleaned
and uncleaned foraminifera are significantly influenced by Nd originating from authigenic Fe-Mn
coatings that precipitate on the foraminiferal shells (Roberts et al., 2012; Tachikawa et al., 2014;
Wu et al., 2015; Yu et al., 2018). Therefore, the Nd isotope composition of both reductively
cleaned and non-reductively cleaned foraminifera reflects the εNd of bottom and/or pore water
and can be used as a useful tracer to reconstruct past seawater εNd changes of bottom water masses.
Accordingly, the foraminiferal εNd records of core LC21 and ODP site 967C should be re-
interpreted as past changes of bottom water conditions and re-evaluated in the light of the new
εNd record presented in this study (Fig. 7).
<figure><img src="image_8.png" /><figcaption>**Fig. 7** Comparison of planktonic foraminiferal $\delta^{18}$O with the $\varepsilon$Nd record during sapropel S5 from four cores from the EMS and two cores from the WMS; (a) northern Levantine Basin: foraminiferal $\varepsilon$Nd (red line, Osborne et al., 2010) and $\delta^{18}$O$_G. ruber$ (blue line, Rohling et al., 2002) of core LC21; foraminiferal $\varepsilon$Nd (red line) and $\delta^{18}$O$_G. ruber$ (blue line) of core MD90-920 (this study); (b) southern Levantine Basin: foraminiferal $\varepsilon$Nd (red line) and $\delta^{18}$O$_G. ruber$ (blue line) of core MD90-964 (Duhamel et al., 2020); foraminiferal $\varepsilon$Nd (red line) and $\delta^{18}$O$_G. ruber$ (blue line) of ODP site 967C, (Scrivner et al., 2004); (c) Tyrrhenian Basin: foraminiferal $\varepsilon$Nd (red line, Colin et al., 2021) and $\delta^{18}$O$_G. bulloides$ (blue line, Toucanne et al., 2012) of core MD01-2472; foraminiferal $\varepsilon$Nd (red line) and $\delta^{18}$O$_G. bulloides$ (blue line) of core MD90-912 (this study). The dashed lines for the core MD90-912 and MD01-2472 are an equivalent time of sapropel deposits. The vertical orange and blue bars are the early and late stage of S5, respectively.</figcaption></figure>
The eastern and western sectors of the Levantine Basin exhibit a general pattern of inverse variations in foraminiferal $\epsilon$Nd (Figs. 7a and 7b). Cores MD90-964 and ODP site 967C, from the eastern Levantine Basin, display an increase in foraminiferal $\epsilon$Nd during the early-stage of the sapropel S5 deposition, accompanied by a decrease in *G. ruber* $\delta^{18}$O values corresponding to a significant decrease of the sea surface salinity (see supplementary text) (Fig. 7b). $\epsilon$Nd values reaches its maximum at ~127 ka for the two southern sites, then decreased until ~123 ka, concurrent with an increase in *G. ruber* $\delta^{18}$O. Finally, the late phase of sapropel S5 (from ~123 to ~120 ka) is characterized by lower foraminiferal $\epsilon$Nd values. In contrast, cores MD90-920 and LC21, located south of the Aegean Sea, show lower $\epsilon$Nd values during the early phase of sapropel S5 (between ~128 and ~123 ka), with values reaching their maximum ($\epsilon$Nd ~-2) towards the end of sapropel S5 (~120-121 ka) (Fig. 7a). Additionally, a small peak in $\epsilon$Nd is observed during the early phase of S5 (~126 ka).

Several studies have documented an intensification of precipitation during the early phase of S5 (from 128 to 123 ka) across two distinct source regions: the northern border of the EMS (e.g., low $\delta^{18}$O values at Soreq cave, Figs. 6b and 8b) and North Africa (Fig. 6b and 8a) through the northward migration of the InterTropical Convergence Zone (ITCZ) during the AHP5 from ~132 to ~126 ka (Amies et al., 2019; Bastian et al., 2017, Ehrmann et al., 2016; Shanahan et al., 2015). This led to significant freshwater discharge into the Levantine Basin, accompanied by water mass stratification (Grant et al., 2012; Schmiedl et al., 2003). Our findings suggest that this also reduced deep-water formation in the Aegean Sea, as highlighted by the low deep current velocities (low Zr/Rb) and anoxic conditions (high S/Al) in core MD90-920 (Fig. 4 and Fig. S1).

As a result, the deep-water circulation of the EMS was significantly reduced and isolated from intermediate and surface layers (Rohling et al., 2015; Tachikawa et al., 2015; Wu et al., 2019). Accordingly, we hypothesize that the stratification of the water column and isolation of the deep-water masses in the Levantine Basin induced a strong local influence on the lithogenic Nd contributions. The increased runoff from the Nile River and other northern African paleo-river systems carried large amounts of freshwater and nutrients to the southern Levantine Basin, resulting in enhanced productivity in the surface water, stagnation in the deep-sea basin, and deposition of organic material (Emeis et al., 2003b; Rossignol-Strick et al., 1982; Rohling, 1994;
Rohling et al., 2015). Sediments from northern African regions are characterized by a high concentration of smectite (from 20 to 40%, up to 85%; Hamann et al., 2009; Ukstins et al., 2002) and a radiogenic εNd signature (from -3 to +3, Tachikawa et al., 2004). Sediments rich in smectite have been shown to exchange Nd efficiently with intermediate- and deep-water of the ocean (Huang et al., 2023, 2024). Therefore, we argue that the input of volcanic material from the Nile and paleo-rivers during the AHP5 (e.g., Revel et al., 2010, 2014; Zhao et al., 2011) and the interaction with the radiogenic sediments of the Nile River deep-sea fan and Egyptian/Libyan margin of the Levantine Basin (Ayache et al., 2016), resulted in a gradual acquisition of radiogenic εNd by the water masses. In addition, the slowdown of the circulation pattern is accompanied by an increase in the residence time of deep-water masses throughout the Levantine Basin (De Lange et al., 2008; Wu et al., 2019). This, in turn, results in a longer contact time between the deep-water masses and the radiogenic sediments, specifically at the radiogenic margin of the Levantine Basin. These processes are in good agreement with the general increase in εNd values as observed in the southern core MD90-964 and ODP Site 967 (Fig. 7b).
<figure><img src="image_9.png" /><figcaption>**Fig.8** (a) The monsoon intensification factors from Amies et al. (2019) and the TOC of core LC21 (Marino et al., 2007); (b) δ18O values of speleothem at Soreq Cave (Central Israel; Bar-Matthews et al., 2003) and δ18O G. ruber of MD90-920 (this study); (c) εNd obtained on mixed planktonic</figcaption></figure>
foraminifera of core MD90-920 (this study) and XRF Zr/Rb ratio of MD90-920 (this study); (d) January temperature of Mediterranean Sea based on pollen (Salonen et al., 2021) and variations in the Relative Sea Level (RSL) (Grant et al., 2012); (e) Summer insolation curve calculated at 15°N in June (Lisiecki et al., 2005) and pollen percentage data for sclerophyllous vegetation at Tenaghi Philippon (Milner et al., 2012). The vertical orange and blue bars are the early and late stage of S5, respectively.

In contrast, the εNd records of the northwestern Levantine Basin (cores MD90-920 and LC21) display a shift towards less radiogenic values at the early-stage of sapropel S5, suggesting that the radiogenic water masses from the southeastern Levantine Basin did not extend to these northern core sites, and/or that these sites were predominantly influenced by greater inputs of northern unradiogenic Nd lithogenic sources. The onset of the sapropel S5 deposition coincided with strong riverine freshwater discharge into the Aegean Sea and a reduction of AeDW formation, as supported by several studies that reconstructed wetter conditions during sapropel S5 in the northern border regions of the EMS (e.g., Frogley et al., 1999; Grant et al., 2017; Mommersteeg et al., 1995; Rossignol-Strick, 1987; Rohling and Hilgen, 1991; Rossignol-Strick, 1995,1995; Rossignol-Strick and Paterne, 1999; Tzedakis, 1993; Wijmstra et al., 1990). Specifically, pollen records of vegetation on the northern continent adjacent to the Levantine Basin show high abundances of deciduous arboreal taxa (especially deciduous oak) and low occurrences of steppe elements, suggesting increased summer rainfall and moisture availability (Fig. 8e) (e.g., Geraga et al., 2010; Kotthoff et al., 2008; Tzedakis, 1993; Triantaphyllou et al., 2009). Moreover, the comparison of the G. ruber δ¹⁸O record of core MD90-920 with the composite isotopic record from speleothems of the Soreq cave (Israel) (Ayalon et al., 2002; Bar-Matthews et al., 2003) reveals periods characterized by high rainfall (i.e., low δ¹⁸O) in the northern border of the EMS at the times of sapropel deposition (Figs. 6b and 8b). Such increase in precipitation on the northern borders of the EMS is particularly well marked at the beginning of S5 (Figs. 6b and 8b) (Allen and Huntley, 2009; Bar-Matthews et al., 2000; Drysdale et al., 2005; Tzedakis et al., 2003; Toucanne et al., 2015). Furthermore, the humidity proxy record derived from the ODP Site 967 points to a prevailing humid climate in North Africa during this period (Grant et al., 2017) (Fig.
6b). Moreover, Kallel et al., (1997 and 2000) used the planktonic foraminiferal fauna records to show that an enhanced precipitations over the whole Mediterranean Sea and the surrounding continents have been invoked to explain Sea surface salinity (SSS) that was everywhere lower than today during sapropel formation and almost homogeneous over the whole Mediterranean Sea. Such humid conditions in the EMS could have potentially induced water mass stratification in the Aegean Sea, a collapse of the AeDW and unradiogenic lithogenic Nd input (Weldeab et al., 2002). As a result of these hydrological conditions, which increased the residence time of deep-water masses, it is likely that cores MD90-920 and LC21 were more influenced by exchanges between seawater and detrital material during the early phase of S5. This influence involved interactions with unradiogenic sediments from the northern margin of the Levantine Basin (εNd from -5 to -3.5; Ayache et al., 2016) and/or with rivers draining the northern surrounding continent of the Levantine Basin and Aegean Sea (εNd between -10 and -6; Frost et al., 1986) (Fig. 9a).

At the late-S5 stage (from ~123 to ~120 ka), the εNd in cores MD90-920 and LC21 increased to more radiogenic values (~ -2.5; Fig. 7a), pointing to a minor reactivation of the circulation, with radiogenic deep-water flowing northwards from the Levantine Basin to the studied sites. The reactivation of the circulation is supported by the observed increase in the Zr/Rb ratio in core MD90-920 during the late-S5 stage (Fig. 8c) and can be attributed to a reduction in riverine freshwater inputs and a cooling of sea surface temperatures (SST) (corresponding to higher G. ruber δ¹⁸O, Fig. 8b) (Kallel et al., 2000; Salonen et al., 2021; Tachikawa et al., 2015). After the onset of the MIS5e (~126 ka ago), Greenland ice core temperatures reached a maximum of 8±4°C, then cooled gradually due to the declining summer insolation (NEEM community members, 2013). A comparable pattern is observed in the January SST curve from the Mediterranean Sea, with a maximum at about 125 ka, followed by a subsequent decrease (Fig. 8d) (Salonen et al., 2021). Discharge from the Nile River is primarily controlled by the movement of the ITCZ, which is influenced by the orbital precession index (e.g., Box et al., 2011; Hennekam et al., 2014; Rossignol-Strick et al., 1982; Rossignol-Strick, 1985; Rohling and De Rijk, 1999; Scrivner et al., 2004; Revel et al., 2010; Weldeab et al., 2014). Ehrmann et al. (2016) showed that the AHP5 persisted for approximately 6 ka (AHP5: 132-126 ka), which does not encompass the entire duration of the sapropel S5 (120-129 ka; Grant et al., 2016). Moreover, Ameis et al. (2019)
developed a quantitative model of the intensification of the North African monsoon, which suggests that monsoon activity reached its maximum at the beginning of S5 and subsequently declined throughout the entire sapropel stage (Fig. 8a). Consequently, the deposition of the late-S5 sapropel occurred in an oceanographic setting characterized by a decline in SST and reduced freshwater input from North Africa, as well as summer aridification of the northern EMS, as evidenced by pollen records (Fig. 8e). These climatic conditions may have resulted in an increase in the density of sea surface water, thereby promoting the formation of deep waters. This gradual recovery of the circulation may have facilitated the transport and mixing of unradiogenic water masses from the northern EMS to the southern Levantine Basin, and conversely, moved radiogenic water masses from the southern EMS to the Aegean Sea. This is consistent with the observed differences in εNd records between the northern and southern Levantine Basin (Fig. 9b).

<figure><img src="image_10.png" /></figure>

(b) late S5 (123-120 ka)

<figure><img src="image_11.png" /><figcaption>**Fig. 9** Schematic representation of the proposed evolution of hydrological circulation during (a) the early stage of sapropel S5; (b) the late stage of sapropel S5.</figcaption></figure>
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## 5.4 Implication in hydrological exchanges between the Levantine Basin and the Tyrrhenian Sea

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Interestingly, the foraminiferal $\epsilon$Nd values from cores MD90-912 (Sardinia Channel) and MD01-2472 (Corsica Trough) (Colin et al., 2021), which lie along the main pathways of the LIW and upper EMDW from the Tyrrhenian Sea to the Balearic Sea, exhibited no major variations at the onset of sapropel S5. However, they displayed more radiogenic signatures around 124-120 ka, coeval with the late-S5 stage in the EMS, with $\epsilon$Nd values reaching -6.8±0.4 and -7.7±0.1, respectively (Figs. 6e and 6f). These results agree with previous studies that have suggested an increase in radiogenic Nd water masses flowing from the EMS to the WMS during periods of interglacial high sea-level (Colin et al., 2021). Nonetheless, there is a notable temporal discrepancy between the foraminiferal $\epsilon$Nd records in the Tyrrhenian Sea and the variations in sea level, as well as with the onset of the deposition of sapropel S5 in the Levantine Basin (Fig. 6g). Consequently, sea level variations cannot be considered the sole mechanism driving the exchange between the EMS and WMS across the Siculo-Tunisian Strait during sapropel depositions.

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The shallow sill depth of the Siculo-Tunisian Strait (average depth < 400 m) results in the EMS and WMS having separate deep-water circulations, with the Adriatic Sea being the main center of the EMDW formation (Milot et al., 1999; Rohling et al., 2015). However, during the 1990s, the Aegean Sea has also been an important dense water source area for the EMS (Roether et al., 1996). The AeDW is characterized by warmer and saltier water that is denser than the EMDW formed in the Adriatic Sea (Chiggiato et al., 2016). During the intensification of the AeDW, a dense deep-water filling of the eastern Mediterranean Basin occurred at greater depths, resulting in an uplift of the previous EMDW to shallower levels (Theocharis et al., 2002). These specific hydrological conditions are favorable to the westward flow of the radiogenic EMDW through the Siculo-Tunisian Strait into the Tyrrhenian Sea (Roether et al., 2007; Sparnocchia et al., 1999; Theocharis et al., 1999, 2002). Consequently, we propose that the reactivation of deep-water formation in the Aegean Sea during the late-S5 stage was responsible for a westward circulation of water masses with high $\epsilon$Nd, resulting from prolonged stratification. This process
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also contributed to the radiogenic εNd peak observed in cores MD90-912 and MD01-2472 from
the Tyrrhenian Sea.

6. Conclusions

The combined geochemical evidence from cores collected in southeastern Aegean Sea (MD90-920) and in the Sardinia Channel (MD90-912), coupled with previous data, offer significant insights into the paleoceanographic conditions during sapropel deposition intervals in the Eastern Mediterranean Sea (EMS). Several negative excursions in planktonic δ¹⁸O values during these periods suggest substantial freshwater influxes, primarily driven by monsoon-fueled discharges from the North African continent. These freshwater inputs led to an intensification of the primary productivity as evidenced by elevated Ba/Al ratio and bottom oxygen depleted conditions, as evidenced by elevated S/Al ratio.

Based on planktonic foraminiferal εNd data, we identified two distinct circulation patterns within the sapropel S5 time interval. In the early stage of S5, a substantial inflow of freshwater induced water stratification, thereby increasing the residence time of bottom water and enhancing boundary exchange processes. This, in turn, imprinted a local εNd signature onto the foraminiferal εNd. Cores collected in the northern Levantine Basin show a less radiogenic signature, while those retrieved in the southern Levantine basin exhibit a more radiogenic signature. During the late stage of S5, colder SST combined with the reduction of the North African monsoon intensity induced the formation of deep-water masses in the Aegean Sea and a slight resumption of a basin-wide circulation. This enhanced circulation progressively mixed radiogenic EMDW, in contact with volcanic Nile River sediments, with unradiogenic AeDW and AdDW in contact with sediments from the northern margin of the EMS.

Overall, our findings indicate, for the first time, that hydrological conditions varied during the deposition of sapropel S5, suggesting an initial phase of stagnation (~128-123 ka) followed by a gradual recovery of circulation in the EMS since ~123 ka and the full restoration after S5 (~120 ka).

Acknowledgements
The research leading to this paper was funded by the French National Research Agency under the "Investissements d'avenir" program (Grant ANR-11-IDEX-0004-17-EURE-0006), the MEDSENS Project (Grant ANR-19-CE01-0019) and the INSU LEFE-IMAGO PALMEDS Project. We gratefully acknowledge the support provided by Louise Bordier during Nd isotopic composition analyses.

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