Introduction
The Mediterranean Sea is a midlatitude semi-enclosed basin, characterized by
evaporation exceeding precipitation and river runoff, where the inflow of
fresh and relatively warm surface Atlantic water is transformed into saltier
and cooler (i.e., denser) intermediate and deep waters. Several studies have
demonstrated that the Mediterranean thermohaline circulation was highly
sensitive to both the rapid climatic changes propagated into the basin from
high latitudes of the Northern Hemisphere (Cacho et al., 1999, 2000, 2002;
Moreno et al., 2002, 2005; Paterne et al., 1999; Martrat et al., 2004; Sierro
et al., 2005; Frigola et al., 2007, 2008) and orbitally forced modifications
of the eastern Mediterranean freshwater budget mainly driven by monsoonal
river runoff from the subtropics (Rohling et al., 2002, 2004; Bahr et al.,
2015). A link between the intensification of the Mediterranean Outflow Water
(MOW) and the intensity of the Atlantic Meridional Overturning Circulation
(AMOC) was proposed (Cacho et al., 1999, 2000, 2001; Bigg and Wadley, 2001;
Sierro et al., 2005; Voelker et al., 2006) and recently supported by new
geochemical data in sediments of the Gulf of Cádiz (Bahr et al., 2015).
In particular, it has been suggested that the intensity of the MOW and, more
generally, the variations of the thermohaline circulation of the
Mediterranean Sea could play a significant role in triggering a switch from a
weakened to an enhanced state of the AMOC through the injection of saline
Mediterranean waters in the intermediate North Atlantic at times of weak AMOC
(Rogerson et al., 2006; Voelker et al., 2006; Khélifi et al., 2009). The
Mediterranean intermediate waters, notably the Levantine Intermediate Water
(LIW), which represent today up to 80 % in volume of the MOW (Kinder and
Parilla, 1987), are considered an important driver of MOW-derived salt into
the North Atlantic. Furthermore, the LIW also plays a key role in controlling
the deep-sea ventilation of the Mediterranean basin, being strongly involved
in the formation of deep waters in the Aegean Sea, Adriatic Sea, Tyrrhenian
Sea and Gulf of Lions (Millot and Taupier-Letage, 2005). It is hypothesized
that a reduction of intermediate and deep-water formation as a consequence of
surface hydrological changes in the eastern Mediterranean basin acted as a
precondition for the sapropel S1 deposition by limiting the oxygen supply to
the bottom waters (De Lange et al., 2008; Rohling et al., 2015; Tachikawa et
al., 2015). Therefore, it is crucial to gain a more complete understanding of
the variability of the Mediterranean intermediate circulation in the past and
its impact on the MOW outflow and, in general, on the Mediterranean
thermohaline circulation.
Previous studies have mainly focused on the glacial variability of the
deep-water circulation in the western Mediterranean basin (Cacho et al.,
2000, 2006; Sierro et al., 2005; Frigola et al., 2007, 2008). During the Last
Glacial Maximum (LGM), strong deep-water convection took place in the Gulf of
Lions, producing cold, well-ventilated western Mediterranean deep water
(WMDW) (Cacho et al., 2000, 2006; Sierro et al., 2005), while the MOW flowed
at greater depth in the Gulf of Cádiz (Rogerson et al., 2005;
Schönfeld and Zahn, 2000). With the onset of the Termination 1 (T1) at
about 15 ka, the WMDW production declined until the onset of the Holocene
due to the rising sea level, with a relatively weak mode during the Heinrich
Stadial 1 (HS1) and the Younger Dryas (YD) (Sierro et al., 2005; Frigola et
al., 2008), which led to the deposition of the Organic Rich Layer 1 (ORL1;
14.5–8.2 ka BP; Cacho et al., 2002).
Because of the disappearance during the Early Holocene of specific
epibenthic foraminiferal species, such as Cibicidoides spp., which are commonly used for
paleohydrological reconstructions, information about the Holocene
variability of the deep-water circulation in the western Mediterranean is
relatively scarce and is mainly based on grain-size analysis and sediment
geochemistry (e.g., Frigola et al., 2007). These authors have identified four
distinct phases representing different deep-water overturning conditions in
the western Mediterranean basin during the Holocene, as well as centennial-
to millennial-scale abrupt events of overturning reinforcement.
Faunal and stable isotope records from benthic foraminifera located at
intermediate depths in the eastern basin reveal well-ventilated LIW during
the last glacial period and deglaciation (Kuhnt et al., 2008; Schmiedl et
al., 2010). Similarly, a grain-size record obtained from a sediment core
collected within the LIW depth range (∼ 500 m water depth) at the
eastern
Corsica margin also documents enhanced bottom currents during the glacial
period and for specific time intervals of the deglaciation, such as HS1 and
YD (Toucanne et al., 2012). The Early Holocene is characterized by a collapse
of the LIW (Kuhnt et al., 2008; Schmiedl et al., 2010; Toucanne et al., 2012)
synchronous with the sapropel S1 deposition (10.2–6.4 cal ka BP; Mercone
et al., 2000). Proxies for deep-water conditions reveal the occurrence of
episodes of deep-water overturning reinforcement in the eastern Mediterranean
basin at 8.2 ka BP (Rohling et al., 1997, 2015; Kuhnt et al., 2007;
Abu-Zied et al., 2008, Siani et al., 2013; Tachikawa et al; 2015),
responsible for the interruption of the sapropel S1 in the eastern
Mediterranean basin (Mercone et al., 2001; Rohling et al., 2015).
Additional insights into Mediterranean circulation changes may be gained
using radiogenic isotopes, such as neodymium, that represent reliable tracers
for constraining water-mass mixing and sources (Goldstein and Hemming, 2003,
and references therein). It has recently been shown that the neodymium (Nd)
isotopic composition, expressed as εNd = ([(143Nd / 144Nd)sample / (143Nd / 144Nd)CHUR] -1) × 10 000 (CHUR: Chondritic Uniform Reservoir, Jacobsen and
Wasserburg, 1980), of living and fossil scleractinian cold-water coral (CWC) faithfully traces
intermediate and deep-water mass provenance and mixing of the ocean (e.g., van
de Flierdt et al., 2010; Colin et al., 2010; López Correa et al., 2012;
Monterro-Serrano et al., 2011, 2013; Copard et al., 2012). Differently from
the CWC, the εNd composition of fossil planktonic foraminifera
is not related to the ambient seawater at calcification depth but reflects
the bottom and/or pore water εNd due to the presence of
authigenic Fe–Mn coatings precipitated on their carbonate shell after
deposition onto the sediment (Roberts et al., 2010; Elmore et al., 2011;
Piotrowski et al., 2012; Tachikawa et al., 2014; Wu et al., 2015). Therefore,
the εNd composition of planktonic foraminiferal tests can be
used as a useful tracer of deep-water circulation changes in the past,
although the effect of pore water on foraminiferal εNd values
could potentially complicate the interpretation (Tachikawa et al., 2014).
In the Mediterranean Sea, modern seawater εNd values display a
large range from ∼ -11 to ∼ -5 and a clear vertical and
longitudinal gradient, with more radiogenic values encountered in the eastern
basin and typically at intermediate and deeper depths (Spivack and Wasserburg
1988; Henry et al., 1994; Tachikawa et al., 2004; Vance et al., 2004).
Considering this large εNd contrast, εNd recorded
in fossil CWC and planktonic foraminifera from the Mediterranean offers great
potential to trace intermediate and deep-water mass exchange between the two
basins, especially during periods devoid of key epibenthic foraminifera, such
as the sapropel S1 or ORL1 events.
Here, the εNd of planktonic foraminifera from a sediment core
collected in the Balearic Sea and CWC samples from the Alboran Sea and the
Sardinia Channel were investigated to establish past changes of the seawater
εNd at intermediate depths and constrain hydrological variations
of the LIW during the last ∼ 20 kyr. The εNd values have
been combined with stable oxygen (δ18O) and carbon
(δ13C) isotope measurements of benthic (Cibicidoides pachyderma) and planktonic (Globigerina bulloides) foraminifera and
sea-surface temperature estimates by modern analogue technique (MAT). Results
reveal significant εNd variations at intermediate depths in the
western basin interpreted as a drastic reduction of the hydrological
exchanges between the western and eastern Mediterranean Sea and the
subsequent higher proportion of intermediate water produced in the Gulf of
Lions during the time interval corresponding to the sapropel S1 deposition.
Seawater εNd distribution in the Mediterranean Sea
The Atlantic Water (AW) enters the Mediterranean Sea as surface inflow
through the Strait of Gibraltar with an unradiogenic εNd
signature of ∼ -9.7 in the strait (Tachikawa et al., 2004) and
∼ -10.4 in the Alboran Sea (Tachikawa et al., 2004; Spivack and
Wasserburg, 1988) for depths shallower than 50 m. During its eastward
flowing, AW mixes with upwelled Mediterranean Intermediate Water forming the
Modified Atlantic Water (MAW) that spreads within the basin (Millot and
Taupier-Letage, 2005) (Fig. 1). The surface water εNd values
(shallower than 50 m) range from -9.8 to -8.8 in the western
Mediterranean basin (Henry et al., 1994; P. Montagna, personal communication, 2016) and -9.3 to -4.2 in the
eastern basin, with seawater off the Nile delta showing the most radiogenic
values (Tackikawa et al., 2004; Vance et al., 2004; P. Montagna, personal communication, 2016). The surface waters in the
eastern Mediterranean basin become denser due to strong mixing and
evaporation caused by cold and dry air masses flowing over the Cyprus–Rhodes
area in winter and eventually sink, leading to the formation of LIW
(Ovchinnikov, 1984; Lascaratos et al., 1993, 1998; Malanotte-Rizzoli et al.,
1999; Pinardi and Masetti, 2000). The LIW spreads throughout the entire
Mediterranean basin at depths between ∼ 150–200 m and
∼ 600–700 m and is characterized by more radiogenic εNd
values ranging from -7.9 to -4.8 (average value ±1σ:
-6.6 ± 1) in the eastern basin and from -10.4 to -7.58
(-8.7 ± 0.9) in the western basin (Henry et al., 1994; Tachikawa et
al., 2004; Vance et al., 2004; P. Montagna,
personal communication, 2016). The LIW acquires its εNd
signature mainly from the partial dissolution of Nile River particles
(Tachikawa et al., 2004), which have an average isotopic composition of
-3.25 (Weldeab et al., 2002), and the mixing along its path with overlying
and underlying water masses with different εNd signatures. The
LIW finally enters the Atlantic Ocean at intermediate depths through the
Strait of Gibraltar with an average εNd value of
-9.2 ± 0.2 (Tachikawa et al., 2004; P. Montagna, personal communication, 2016).
Map of the western Mediterranean Sea showing the locations of
samples investigated in this study. Yellow dot indicates the sampling
location of the sediment core from the Balearic Sea (SU92-33); yellow stars
indicate the locations of the CWC-bearing cores from the Sardinia Channel
(RECORD 23) and the southern Alboran Sea (for further details on the CWC
from the Alboran Sea refer also to Fink et al., 2013). The cores discussed
in this paper (Gulf of Cádiz: IODP site U1387, Balearic Sea: MD09-2343,
northern Tyrrhenian Sea: MD01-2472, Adriatic Sea: MD90-917) are indicated by
black dots, and seawater stations are marked by open squares. Arrows
represent the main oceanographic currents. The black line shows the general
trajectory of the Modified Atlantic Water (MAW) flowing at the surface from
the Atlantic Ocean toward the western and eastern Mediterranean. The orange
line represents the Levantine Intermediate Water (LIW) originating from the
eastern basin. The black dashed line shows the trajectory of the western
Mediterranean deep water (WMDW) flowing from the Gulf of Lions toward the
Strait of Gibraltar.
The WMDW is formed in the Gulf of Lions due to winter cooling and evaporation
followed by mixing between surface waters and the more saline LIW and spreads
into the Balearic basin and Tyrrhenian Sea between ∼ 2000 and 3000 m
(Millot, 1999; Schroeder et al., 2013) (Fig. 1). The WMDW is characterized by
an average εNd value of -9.4 ± 0.9 (Henry et al., 1994;
Tachikawa et al., 2004; P. Montagna, personal
communication, 2016). Between the WMDW and the LIW (from ∼ 700 to
2000 m), the Tyrrhenian deep water (TDW) has been found (Millot et al.,
2006), which is produced by the mixing between WMDW and eastern Mediterranean
deep water (EMDW) that cascades in the Tyrrhenian Sea after entering through
the Strait of Sicily (Millot, 1999, 2009; Astraldi et al., 2001). The TDW has
an average εNd value of -8.1 ± 0.5 (P. Montagna, personal communication, 2016).
Results
Cold-water corals
The good state of preservation for the CWC samples from the Sardinia Channel
(RECORD 23; Fig. 1) is attested by their initial δ234U values
(Table 1), which is in the range of the modern seawater value
(146.8 ± 0.1; Andersen et al., 2010). If the uncertainty of the δ234Ui is taken into account, all the values fulfill the so-called
“strict” ±4 ‰ reliability criterion and the U / Th ages
can be considered strictly reliable. The coral ages range from
0.091 ± 0.011 to 10.904 ± 0.042 ka BP (Table 1) and reveal
three distinct clusters of coral age distribution during the Holocene that
represent periods of sustained coral occurrence. These periods coincide
with the Early Holocene encompassing a 700-year time interval from
∼ 10.9 to 10.2 ka BP, the very late Early Holocene at
∼ 8.7 ka BP, and the Late Holocene starting at ∼ 1.5 ka BP
(Table 1).
(a) Sea-surface temperature (SST) record of core SU92-33 (red
line), (b) εNd records obtained on mixed planktonic
foraminifers from core SU92-33 (open circles) and from cold-water coral
fragments collected in the Alboran Sea (red squares) and (c) εNd
values of cold-water corals from core RECORD 23 (Sardinia Channel).
Radiocarbon ages obtained for CWC samples collected in the Alboran Sea were
published by Fink et al. (2013) (Table 2). They also document three periods
of sustained CWC occurrence coinciding with the Bølling–Allerød (B-A)
interstadial (13.5–12.9 cal ka BP), the Early Holocene
(11.2–9.8 cal ka BP) and the Mid- to Late Holocene
(5.4–0.3 cal ka BP).
The εNd record obtained from the CWC samples from the Alboran
Sea displays a narrow range from -9.22 ± 0.30 to -8.59 ± 0.3,
which is comparable to the εNd record of the planktonic
foraminifera from the Balearic Sea over the last 13.5 kyr (Table 2,
Fig. 3b). Most of the CWC εNd values are similar within the
analytical error and the record does not reveal any clear difference over the
last ∼ 13.5 kyr.
On the contrary, the CWC samples from the Sardinia Channel display a
relatively large εNd range, with values varying from
-5.99 ± 0.50 to -7.75 ± 0.10 during the Early and Late
Holocene and values as low as -8.66 ± 0.30 during the
mid-sapropel S1 deposition (S1a) at ∼ 8.7 ka BP (Table 1, Fig. 3c).
Core SU92-33
The stratigraphy of core SU92-33 was derived from the δ18O
variations of the planktonic foraminifera G. bulloides (Fig. 2b).
The last glacial–interglacial transition and the Holocene encompasses the
upper 2.1 m of the core (Fig. 2b). The δ18O record of G.
bulloides shows higher values (∼ 3.5 ‰) during the
late glacial compared to the Holocene (from ∼ 1.5 to 0.8 ‰),
exhibiting a pattern similar to those observed in nearby deep-sea cores from
the western Mediterranean Sea (Sierro et al., 2005; Melki et al., 2009).
The age model for the upper 1.2 m of the core SU92-33 was based on seven
AMS 14C age measurements and a linear interpolation between these ages
(Table 3, Fig. 2). For the lower portion of the core, a control point was
established at the onset of the last deglaciation, which is coeval in the
western and central Mediterranean seas at ∼ 17 cal ka BP (Sierro et
al., 2005; Melki et al., 2009; Siani et al., 2001). Overall, the upper 2.1 m
of core SU92-33 span the last 19 kyr, with an estimated average
sedimentation rate ranging from ∼ 15 cm ka-1 during the
deglaciation to ∼ 10 cm ka-1 during the Holocene.
April–May SST reconstruction was derived from MAT to define the main climatic
events recorded in core SU92-33 during the last 19 kyr. SSTs vary from 8.5 to
17.5 ∘C with high amplitude variability over the last 19 kyr BP
(Fig. 2a). The LGM (19–18 ka BP) is characterized by SST values centered
at around 12 ∘C. Then, a progressive decrease of
∼ 4 ∘C between 17.8 and 16 ka marks the HS1 (Fig. 2a). A warming phase (∼ 14 ∘C) between 14.5 and
13.8 ka BP coincides with the B-A interstadial and is followed by a cooling
(∼ 11 ∘C) between 13.1 and 11.8 ka BP largely corresponding
to the YD (Fig. 2a). During the Holocene, SSTs show mainly values of
∼ 16 ∘C, with one exception between 7 and 6 ka BP pointing
to an abrupt cooling of ∼ 3 ∘C (Fig. 2a). From the late
glacial to the Holocene, SST variations show a similar pattern to that
previously observed in the Gulf of Lions and Tyrrhenian Sea (Kallel et al.,
1997; Melki et al., 2009) as well as in the Alboran Sea (Martrat et al.,
2014; Rodrigo-Gámiz et al., 2014). They are globally synchronous for the
main climatic transitions to the well-dated south Adriatic Sea core MD90-917
(Siani et al., 2004) confirming the robustness of the SU92-33 age model
(Fig. 2a).
The δ18O and δ13C records obtained from the benthic
foraminifera C. pachyderma display significant variations at
millennial timescales (Fig. 2c, d). The δ18O values decrease
steadily from ∼ 4.5 ‰ during the LGM to ∼ 1.5 ‰
during the Holocene, without showing any significant excursion during HS1 and
the YD events (Fig. 2c), in agreement with results obtained from the neighbor
core MD99-2343 (Sierro et al., 2005).
The δ13C record of C. pachyderma shows a decreasing trend
since the LGM with a low variability from ∼ 1.6 to
∼ 0.6 ‰ (Fig. 2d). The heaviest δ13C values are
related to the LGM (∼ 1.6 ‰) while the lightest values
(∼ 0.6 ‰) characterize the Early Holocene and in particular
the period corresponding to the sapropel S1 event in the eastern
Mediterranean basin (Fig. 2d).
The εNd values of planktonic foraminifera of core SU92-33 from
the Balearic Sea vary within a relatively narrow range between
-9.50 ± 0.30 and -8.61 ± 0.30, with an average value of
-9.06 ± 0.28 (Table 2, Fig. 3b). The record shows a slight increasing
trend since the LGM, with the more unradiogenic values (average
-9.28 ± 0.15; n= 7) being observed in the oldest part of the
record (between 18 and 13.5 ka BP), whereas Holocene values are generally
more radiogenic (average -8.84 ± 0.22; n= 17) (Fig. 3b).
Discussion
Overall, the CWC and foraminiferal εNd values measured in this
study point to a pronounced dispersion at intermediate depth in terms of
absolute values and variability in Nd isotopes during the Holocene between
the Alboran and Balearic seas and the Sardinia Channel. Furthermore, the
foraminiferal εNd record reveals an evolution towards more
radiogenic values at intermediate water depth in the Balearic Sea over the
last ∼ 19 kyr (Fig. 3).
A prerequisite to properly interpret such εNd differences and
variations through time consists in characterizing first the present-day
εNd of the main water-mass endmembers present in the western
Mediterranean basin. It is also necessary to evaluate the temporal changes
in εNd of the endmembers since the LGM and assess the
potential influences of lithogenic Nd input and regional exchange between
the continental margins and seawater (“boundary exchange”; Lacan and
Jeandel, 2001, 2005) on the εNd values of intermediate water
masses.
During its westward flow, the LIW continuously mixes with surrounding waters
with different εNd signatures lying above and below. For the
western Mediterranean basin, these water masses are the MAW–Western
Intermediate Water (WIW) and the TDW–WMDW. As a result, a gradual εNd gradient exists at intermediate depth between the eastern and western
Mediterranean basins, with LIW values becoming progressively more
unradiogenic towards the Strait of Gibraltar, from -4.8 ± 0.2 at
227 m in the Levantine basin to -10.4 ± 0.2 at 200 m in the Alboran
Sea (Tachikawa et al., 2004). Such an εNd pattern implies an
effective vertical mixing with more unradiogenetic water masses along the E-W
LIW trajectory ruling out severe isotopic modifications of the LIW due to the
local exchange between the continental margins and seawater. Unfortunately,
no information exists on the potential temporal variability in εNd of the Mediterranean water-mass endmembers since the LGM.
It has been demonstrated that eolian dust input can modify the surface and
subsurface εNd distribution of the ocean in some areas (Arsouze
et al., 2009). The last glacial period was associated with an aridification
of North Africa (Sarnthein et al., 1981; Hooghiemstra et al., 1987; Moreno et
al., 2002; Wienberg et al., 2010) and higher fluxes of Saharan dust to the NE
tropical Atlantic (Itambi et al., 2009) and the western Mediterranean Sea
characterized by unradiogenic εNd values (between
-11 ± 0.4 and -14 ± 0.4; see synthesis in Scheuvens et al.,
2013). Bout-Roumazeilles et al. (2013) documented a dominant role of eolian
supply in the Siculo–Tunisian Strait during the last 20 ka, with the
exception of a significant riverine contribution (from the Nile River) and a
strong reduction of eolian input during the sapropel S1 event. Such
variations in the eolian input to the Mediterranean Sea are not associated to
a significant change in the seawater εNd record obtained for the
Balearic Sea (core SU92-33) during the sapropel S1 event (Fig. 3).
Furthermore, the εNd signature of the CWC from the Sardinia
Channel (core RECORD 23) shifts to more unradiogenic values
(-8.66 ± 0.30) during the sapropel S1 event, which is opposite to
what would be expected from a strong reduction of eolian sediment input. In a
recent study, Rodrigo-Gámiz et al. (2015) have documented variations in
the terrigenous provenance from a sediment record in the Alboran Sea (core
293G; 36∘10.414′ N, 2∘45.280′ W; 1840 m water depth)
since the LGM. Radiogenic isotopes (Sr, Nd, Pb) point to changes from North
African dominated sources during the glacial period to European dominated
source during the Holocene. Nevertheless, the major Sr–Nd–Pb excursions
documented by Rodrigo-Gámiz et al. (2015) and dated at ca. 11.5, 10.2,
8.9–8.7, 5.6, 2.2 and 1.1 ka cal BP do not seem to affect the
εNd values of our foraminifera and coral records.
Taken together, these results suggest that changes of eolian dust input
since the LGM cannot explain the observed εNd variability at
intermediate water depths.
Consequently, assuming that the Nd isotopic budget of the western
Mediterranean Sea has not been strongly modified since the LGM, the
reconstructed variations of the E-W gradient of εNd values in
the western Mediterranean Sea for the past and notably during the sapropel S1
event (Fig. 3) are indicative of a major reorganization of intermediate water
circulation.
(a) δ13C records obtained on benthic foraminifer
C. pachyderma for cores SU92-33 (red line) and MD99-2343 (blue line;
Sierro et al., 2005). (b) εNd records obtained on mixed
planktonic foraminifers from core SU92-33 (open circles) and from cold-water
coral fragments collected in the Alboran Sea (red squares). Modern
εNd values for LIW (orange dashed line) and WMDW (blue dashed
line) are also reported for comparison. (c) εNd values
obtained for planktonic foraminifera with Fe–Mn coatings at sites 300G
(36∘21.532′ N, 1∘47.507′ W; 1860 m; open dots) and
304G (36∘19.873′ N, 1∘31.631′ W; 2382 m; black dots)
in Alboran Sea (Jimenez-Espejo et al., 2015). (d) UP10 fraction
(> 10 µm) from core MD99-2343 (Frigola et al., 2008).
(e) Sortable silt mean grain size of core MD01-2472 (Toucanne et
al., 2012). (f) Ln Zr / Al ratio at IODP site U1387
(36∘48.3′ N 7∘43.1′ W; 559 m) (Bahr et al., 2015).
Hydrological changes in the Alboran and Balearic seas since the LGM
The range in εNd for the CWC from the Alboran Sea (from
-9.22 ± 0.30 to -8.8.59 ± 0.30; Table 2) is very close to the
one obtained for the planktonic foraminifera from the Balearic Sea (from
-9.50 ± 0.30 to -8.61 ± 0.30; Table 4, Fig. 3c), suggesting
that both sites are influenced by the same intermediate water masses at least
for the last 13.5 kyr BP. Today, LIW occupies a depth range between
∼ 200 and ∼ 700 m in the western Mediterranean basin (Millot,
1999; Sparnocchia et al., 1999). More specifically, the salinity maximum
corresponding to the core of LIW is found at around 400 m in the Alboran Sea
(Millot, 2009) and up to 550 m in the Balearic Sea (López-Jurado et al.,
2008). The youngest CWC sample collected in the Alboran Sea with a rather
“recent” age of 0.34 cal ka BP (Fink et al., 2013) displays an
εNd value of -8.59 ± 0.30 (Table 2) that is similar to
the present-day value of the LIW at the same site (-8.3 ± 0.2)
(Dubois-Dauphin et al., 2016) and is significantly different from the
WMDW εNd signature in the Alboran Sea (-10.7 ± 0.2,
1270 m water depth; Tachikawa et al., 2004). Considering the intermediate
depth range of the studied CWC and foraminifera samples, we can reasonably
assume that samples from both sites, in the Balearic Sea (622 m water depth)
and in the Alboran Sea (280 to 442 m water depth), record εNd
variations of the LIW. The εNd record obtained from planktonic
foraminifera generally displays more unradiogenic and homogenous values
before ∼ 13 cal ka BP (range from -9.46 to -9.12) compared to
the most recent part of the record (range from -9.50 to -8.61), with the
highest value of -8.61 ± 0.3 in the Early and Late Holocene.
The SST record displays values centered at around 12 ∘C during the
LGM with a subsequent rapid SST decrease towards 9 ∘C, highlighting
the onset of the HS1 (Fig. 2a). These values are comparable to recent
high-resolution SST data obtained in the Alboran Sea (Martrat et al., 2014;
Rodrigo-Gámiz et al., 2014).
The δ18O record obtained on G. bulloides indicates an
abrupt 1 ‰ excursion towards lighter values centered at about
16 cal ka BP (Table 4), synchronous with the HS1 (Fig. 2b), which is
similar to the δ18O shift reported by Sierro et al. (2005) for a
core collected at 2391 m water depth NE of the Balearic Islands (MD99-2343;
Fig. 1). As the Heinrich events over the last glacial period are
characterized by colder and fresher surface water in the Alboran Sea (Cacho
et al., 1999; Pérez-Folgado et al., 2003; Martrat et al., 2004, 2014;
Rodrigo-Gámiz et al., 2014) and dry climate on land over the western
Mediterranean Sea (Allen et al., 1999; Combourieu-Nebout et al., 2002;
Sanchez Goni et al., 2002; Bartov et al., 2003), lighter δ18O
values of planktonic G. bulloides are thought to be the result of
the inflow of freshwater derived from the melting of icebergs in the Atlantic
Ocean into the Mediterranean Sea (Sierro et al., 2005; Rogerson et al.,
2008).
During this time interval, the δ13C record of C. pachyderma from the Balearic Sea (core SU92-33) displays a decreasing
δ13C trend after ∼ 16 cal ka BP (from 1.4 to
0.9 ‰; Table 4; Fig. 4a). Moreover, the δ13C record
obtained on benthic foraminifera C. pachyderma from the deep
Balearic Sea (core MD99-2343) reveals similar δ13C values before
∼ 16 cal ka BP, suggesting well-mixed and ventilated water masses
during the LGM and the onset of the deglaciation (Sierro et al., 2005).
The slightly lower foraminiferal εNd values before
∼ 13 cal ka BP could reflect a stronger influence of water masses
deriving from the Gulf of Lions as WMDW (εNd:
-9.4 ± 0.9; Henry et al., 1994; Tachikawa et al., 2004;
P. Montagna, personal communication, 2016).
This is in agreement with εNd results obtained by
Jiménez-Espejo et al. (2015) from planktonic foraminifera collected from
deep-water sites (1989 and 2382 m) in the Alboran Sea (Fig. 4c).
Jiménez-Espejo et al. (2015) documented lower εNd values
(ranging from -10.14 ± 0.27 to -9.58 ± 0.22) during the LGM,
suggesting an intense deep-water formation. This is also associated to an
enhanced activity of the deeper branch of the MOW in the Gulf of Cádiz
(Rogerson et al., 2005; Voelker et al., 2006) linked to the active production
of the WMDW in the Gulf of Lions during the LGM (Jiménez-Espejo et al.,
2015).
The end of the HS1 (14.7 cal ka BP) is concurrent with the onset of the
B-A warm interval characterized by increased SSTs up to 14 ∘C in the
Balearic Sea (SU92-33: Fig. 3a), also identified for various sites in the
Mediterranean Sea (Cacho et al., 1999; Martrat et al., 2004, 2014; Essallami
et al., 2007; Rodrigo-Gámiz et al., 2014). The B-A interval is associated
with the so-called meltwater pulse 1A (e.g., Weaver et al., 2003) occurring
at around 14.5 cal ka BP. This led to a rapid sea-level rise of about
20 m in less than 500 years and large freshwater discharges in the Atlantic
Ocean due to the melting of continental ice sheets (Deschamps et al., 2012),
resulting in an enhanced Atlantic inflow across the Strait of Gibraltar.
Synchronously, cosmogenic dating of Alpine glacier retreat throughout the
western Mediterranean hinterland suggests maximum retreat rates (Ivy-Ochs et
al., 2007; Kelly et al., 2006). Overall, these events are responsible for
freshening Mediterranean waters and reduced surface water density and, hence,
weakened ventilation of intermediate (Toucanne et al., 2012) and deep-water
masses (Cacho et al., 2000; Sierro et al., 2005). Similarly, lower benthic
δ13C values obtained for the Balearic Sea (Fig. 4a) point to less
ventilated intermediate water relative to the late glacial. In addition, a
decoupling in the benthic δ13C values is observed between deep
(MD99-2343) and intermediate (core SU92-33) waters after
∼ 16 cal ka BP (Sierro et al. 2005), suggesting an enhanced
stratification of the water masses (Fig. 4a). At this time, the shallowest
εNd record from the deep Alboran Sea (core 300G) shifted towards
more radiogenic values, while the deepest one (core 304G) remained close to
the LGM values (Jimenez-Espejo et al., 2015) (Fig. 4c). Furthermore, results
from the UP10 fraction (particles > 10 µm) of the
MD99-2343 sediment core (Fig. 4d) indicate a declining bottom-current
velocity at 15 ka BP (Frigola et al., 2008). Rogerson et al. (2008) have
hypothesized that during deglacial periods the sinking depth of dense waters
produced in the Gulf of Lions was shallower resulting in new intermediate
water (WIW) rather than new deep water (WMDW) as observed today during mild
winters (Millot, 1999; Schott et al., 1996). Therefore, intermediate depths
of the Balearic Sea could have been isolated from the deep water with the
onset of the T1 (at ∼ 15 ka BP). The reduced convection in the deep
western Mediterranean Sea together with the shoaling of the nutricline
(Rogerson et al., 2008) led to the deposition of the ORL 1 (14.5 to
8.2 ka BP; Cacho et al., 2002) and dysoxic conditions below 2000 m in
agreement with the absence of epibenthic foraminifera such as C. pachyderma after 11 cal ka BP in MD99-2343 (Sierro et al., 2005)
(Fig. 4a).
(a) δ18O record obtained on planktonic foraminifer
G. bulloides for core SU92-33, (b) δ13C records obtained on benthic
foraminifer C. pachyderma for core SU92-33, (c) εNd values of cold-water
corals from core RECORD 23 (Sardinia Channel), (d) εNd values
records obtained on mixed planktonic foraminifera from core SU92-33 (open
circles) and from cold-water coral fragments collected in the Alboran Sea
(red squares) and (e) εNd values obtained on terrigenous fraction
of MS27PT located close the Nile River mouth in the eastern Mediterranean
basin (Revel et al., 2015).
After 13.5 ka BP, planktonic foraminifera εNd values from the
Balearic Sea (core SU92-33) become more radiogenic and are in the range of
CWC εNd values from the Alboran Sea (Fig. 4b). These values may
reveal a stronger influence of the LIW in the Balearic Sea during the YD, as also supported by the sortable silt record from the Tyrrhenian Sea
(Toucanne et al., 2012) (Fig. 4e). Deeper depths of the Alboran Sea also
record a stronger influence of the LIW with an εNd value of
-9.1 ± 0.4 (Jimenez-Espejo et al., 2015). In addition, a concomitant
activation of the upper MOW branch, as reconstructed from higher values of
Zr / Al ratio in sediments of the Gulf of Cádiz, can be related to the
enhanced LIW flow in the western Mediterranean Sea (Fig. 4f) (Bahr et al.,
2015).
The time of sapropel S1 deposition (10.2–6.4 ka) is characterized by a
weakening or a shutdown of intermediate- and deep-water formation in the
eastern Mediterranean basin (Rossignol-Strick et al., 1982; Cramp and
O'Sullivan, 1999; Emeis et al., 2000; Rohling et al., 2015). At this time,
planktonic foraminifera εNd values from intermediate water
depths in the Balearic Sea (core SU92-33) remain high (between
-9.15 ± 0.3 and -8.61 ± 0.3) (Fig. 4b). In contrast,
the deeper Alboran Sea provides a value of -9.8 ± 0.3 pointing to a
stronger contribution of WMDW (Jimenez-Espejo et al., 2015), coeval with the
recovery of deep-water activity from core MD99-2343 (Frigola et al., 2008).
Hydrological changes in the Sardinia Channel during the Holocene
The present-day hydrographic structure of the Sardinia Channel is
characterized by four water masses, with the surface, intermediate and
deep-water masses being represented by MAW, LIW and TDW–WMDW, respectively
(Astraldi et al., 2002a; Millot and Taupier-Lepage, 2005). In addition, the
WIW, flowing between the MAW and the LIW, has also been observed along the
Channel (Sammari et al., 1999). The core of the LIW is located at 400–450 m
water depth in the Tyrrhenian Sea (Hopkins, 1988; Astraldi et al., 2002b),
which is the depth range of CWC samples from the Sardinia Channel (RECORD 23;
414 m) (Taviani et al., 2015). The youngest CWC sample dated at
∼ 0.1 ka BP has an εNd value of -7.70 ± 0.10
(Table 1, Fig. 5), which is similar within error to the value obtained from a
seawater sample collected at 451 m close to the coral sampling location
(-8.0 ± 0.4; P. Montagna, personal
communication, 2016).
The CWC dating from the Sardinia Channel shows three distinct periods of
sustained coral occurrence in this area during the Holocene, with each
displaying a large variability in εNd values. CWC from the Early
Holocene (10.9–10.2 ka BP) and the Late Holocene (< 1.5 ka BP)
exhibit similar ranges of εNd values (ranging from
-5.99 ± 0.50 to -7.75 ± 0.20; Table 1, Fig. 5c). Such
variations are within the present-day εNd range being
characteristic for intermediate waters in the eastern Mediterranean Sea
(-6.6 ± 1.0; Tachikawa et al., 2004; Vance et al., 2004). However, the
CWC εNd values are more radiogenic than those observed at
mid-depth in the present-day western basin (ranging from -10.4 ± 0.2
to -7.58 ± 0.47; Henry et al., 1994; Tachikawa et al., 2004; P.
Montagna, personal communication, 2016),
suggesting a stronger LIW component in the Sardinia Channel during the Early
and Late Holocene. The Sardinian CWC εNd variability also
reflects the sensitivity of the LIW to changes in the eastern basin such as
rapid variability of the Nile River flood discharge (Revel et al., 2014,
2015; Weldeab et al., 2014) or a modification through time in the proportion
between the LIW and the Cretan Intermediate Water (CIW). Today, the
intermediate water outflowing from the Strait of Sicily is composed by
∼ 66 to 75 % of LIW and 33 to 25 % of CIW (Manca et al., 2006;
Millot, 2014). As the CIW is formed in the Aegean Sea, this intermediate
water mass is generally more radiogenic than LIW (Tachikawa et al., 2004; P.
Montagna, personal communication, 2016).
Following this hypothesis, a modification of the mixing proportion between
the CIW and the LIW may potentially explain values as radiogenic as about -6
in the Sardinia Channel during the Early and Late Holocene (Fig. 5c).
However, a stronger LIW and/or a CIW contribution cannot be responsible for
εNd values as low as -8.66 ± 0.30 observed during the
sapropel S1 event at 8.7 ka BP (Table 1, Fig. 5c). Considering that such
unradiogenic value is not observed at intermediate depth in the modern
eastern Mediterranean basin, the most plausible hypothesis suggested here is
that the CWC were influenced by a higher contribution of intermediate water
from the western basin.
Hydrological implications for the intermediate water masses of the western Mediterranean Sea
The εNd records of the Balearic Sea, Alboran Sea and Sardinia
Channel document a temporal variability of the east-west gradient in the
western Mediterranean basin during the Holocene. The magnitude of the
gradient ranges from ∼ 1.5 to ∼ 3 ε units during the
Early and Late Holocene and it is strongly reduced at 8.7 ka BP (from 0 to
∼ 0.5 ε unit), coinciding with the sapropel S1 event
affecting the eastern Mediterranean basin (Fig. 5). Such variations could be
the result of a modification of the Nd isotopic composition of intermediate
water masses due to changes of the LIW production through time and a higher
contribution of the western-sourced intermediate water towards the Sardinia
Channel coinciding with the sapropel S1 event.
The LIW acquires its radiogenic εNd signature in the
Mediterranean Levantine basin mainly from Nd exchange between seawater and
lithogenic particles originating mainly from Nile River (Tachikawa et al.,
2004). A higher sediment supply from the Nile River starting at
∼ 15 ka BP was documented by a shift to more radiogenic εNd values of the terrigenous fraction obtained from a sediment core having
been influenced by the Nile River discharge (Revel et al., 2015) (Fig. 5e).
Other studies pointed to a gradual enhanced Nile River runoff as soon as
14.8 ka BP and a peak of Nile discharge from 9.7 to 8.4 ka recorded by
large increase in sedimentation rate from 9.7 to 8.4 ka
(> 120 cm ka-1) (Revel et al., 2015; Weldeab et al., 2014;
Castaneda et al., 2016). Similarly, enhanced Nile discharge at
∼ 9.5 cal kyr BP was inferred based on δ18O in planktonic
foraminifera from a sediment core in the southeast Levantine Basin (PS009PC; 32∘07.7′ N, 34∘24.4′ E; 552 m water depth) (Hennekam
et al., 2014). This increasing contribution of the Nile River to the eastern
Mediterranean basin has been related to the African Humid Period
(14.8–5.5 ka BP; Shanahan et al., 2015), which in turn was linked to the
precessional increase in Northern Hemisphere insolation during low
eccentricity (deMenocal et al., 2000; Barker et al., 2004; Garcin et al.,
2009). An increasing amount of radiogenic sediments dominated by the
Blue Nile–Atbarah River contribution (Revel et al., 2014) could have modified
the εNd of surface water towards more radiogenic values
(M. Revel, personal communication, 2016).
Indeed, planktonic foraminifera εNd values as high as ∼ -3
have been documented in the eastern Levantine Basin (ODP site 967;
34∘04.27′ N, 32∘43.53′ E; 2553 m water depth) during
the sapropel S1 event as a result of enhanced Nile flooding (Scrivner et al.,
2004). The radiogenic signature was likely transferred to intermediate depth
as a consequence of the LIW formation in the Rhodes Gyre, and it might have
been propagated westwards towards the Sardinia Channel.
Therefore, considering the more unradiogenic value of the CWC samples from
the Sardinia Channel during the sapropel S1a event, it is very unlikely that
eastern-sourced water flowed at intermediate depth towards the Sardinia
Channel. A possible explanation could be the replacement of the radiogenic
LIW that was no longer produced in the eastern basin (Rohling, 1994) by less
radiogenic western intermediate water (possibly WIW). Such a scenario could
even support previous hypotheses of a potential circulation reversal in the
eastern Mediterranean from anti-estuarine to estuarine during sapropel
formation (Huang and Stanley, 1972; Calvert, 1983; Sarmiento et al., 1988;
Buckley and Johnson, 1988; Thunell and Williams, 1989). An alternative
hypothesis would be that reduced surface water densities in the eastern
Mediterranean during sapropel S1 resulted in the LIW sinking to shallower
depths than at present. In this case, CWC from the Sardinia Channel would
have been bathed by underlying WIW during the
sapropel S1a event.