Marine Geology 375 (2016) 120–133
Contents lists available at ScienceDirect
Marine Geology
journal homepage: www.elsevier.com/locate/margo
The 1987 Aegean dense water formation: A streamtube investigation by
comparing theoretical model results, satellite, field, and numerical data
with contourite distribution
M. Bellacicco a,b, C. Anagnostou c, F. Falcini a, E. Rinaldi a, K. Tripsanas d, E. Salusti a,⁎
a
Istituto di Scienze dell'Atmosfera e del Clima (ISAC)-CNR, Via Fosso del Cavaliere 100, Roma, Italy
Università degli Studi di Napoli “Parthenope”, Via Ammiraglio Acton, 38, Napoli, Italy
c
Hellenic Centre for Marine Research, Institute of Oceanography, 46.7 km Athens-Sounio Av., 19013 Anavyssos, Greece
d
Geology Home Team, Shell U.K. Limited, Aberdeen, AB 12 3FY, United Kingdom
b
a r t i c l e
i n f o
Article history:
Received 16 December 2014
Received in revised form 13 January 2016
Accepted 25 January 2016
Available online 27 January 2016
Keywords:
EMT
Aegean Sea
Dense water formation
Contourites
a b s t r a c t
We here discuss a detailed investigation of the dense water formation, evolution and spreading in the Aegean Sea
during the year 1987, immediately prior to the onset of the Eastern Mediterranean Transient (EMT). We use hydrologic data collected during the LIA cruise; satellite images for SST (Sea Surface Temperature), and PROTHEUS
data (a coupled ocean–atmosphere numeric model) along with theoretical streamtube models. These hydrological analyses are related to late Quaternary sedimentary drifts in the Cyclades Plateau and in the Myrtoon Basin.
Our analysis shows that streamtube dynamics provide a novel model of dense water evolution and spreading in
the Aegean Basin. Applying this model to dense water masses observed in winter and spring 1987 near Samothrace and over the Limnos-Lesbos Plateau, results in a geostrophic flow of this dense, cold water towards the
Limnos-Sporades Channel, in the North Aegean Sea. There it mixes with dense water from the Limnos-Lesbos Plateau and finally both move geostrophically towards the Cyclades Plateau. These results indicate that most of the
dense water observed near the Cyclades, formed initially about 3 months earlier at Samothrace and Limnos
shelves. During its long pathway it partially mixed with adjacent water masses. Although our analysis concerns
only one year of dense water analyses, these results are thought to reflect a more general and recurrent phenomenon in the Aegean basin. Indeed, high-resolution (Airgun 10 in.) seismic-reflection data from the Cyclades Plateau reveal the presence of late Quaternary sediment drifts. These observations are concordant with results from
our theoretical model. This suggests a direct link between such a dense-water cascading and contourite dynamics. The continuation of sediment drifts into the deep basin floor (≈900 m deep) of the Myrtoon Basin, moreover,
indicates a cascading character of such bottom currents at the flanks of the basin, a feature that set further
investigations.
© 2016 Elsevier B.V. All rights reserved.
1. Introduction
The idea that the Aegean Sea could be the source of the Eastern Mediterranean deep waters was first proposed by Nielsen (1912), but successive observations demonstrated the main role of the Adriatic Sea as
source of such bottom waters (Malanotte-Rizzoli and Hecht, 1988, and
references therein). Field observations from the Aegean Sea around
the year 1990 revived the Nielsen hypothesis by demonstrating that
the Cretan dense water (CDW) was observed on the Northern Aegean
coasts and on the shallow shelves of the Cyclades islands, and then accumulated in the Cretan Sea (Zervakis et al., 2000; Gertman et al.,
2006). CDW, later on, filled the bottom layers of both Mediterranean
Sea basins (Roether et al., 1996; Klein et al., 1999) although, overflowing
⁎ Corresponding author.
E-mail address: [email protected] (E. Salusti).
http://dx.doi.org/10.1016/j.margeo.2016.01.012
0025-3227/© 2016 Elsevier B.V. All rights reserved.
the Sicily Strait, this deep water results to be heavily modified (Ben
Ismail et al., 2014).
During 1987, before the EMT event, an unexpected dense water was
observed in the Aegean Sea. Its density (σθ ≈ 29.25–29.35) was indeed
higher than the historical maximum value of 29.2 (Georgopoulos et al.,
1992; Theocharis and Georgopoulos, 1993, two similar articles called
G2T3 in the following).
Whereas the period 1992–1993 is generally considered as the real
onset of a full EMT (Roether et al., 1996), the cold winter of the year
1987 is of some interest since modern field data are available and also
because of a particularly severe 3–13 March storms (Lagouvardos
et al., 1998) due to cold fronts passing over the Aegean Sea.
About the EMT event, during September 1987 the first MeteorPOEM cruise did not find any presence of the CDW outflow into the
deep Eastern Mediterranean, while Gertman et al. (2006) evidenced a
CDW outflow from the Island of Kassos in the Cretan Arc (Fig. 1) during
M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
121
Fig. 1. General map of the Aegean Sea bathymetry, and the hydrologic transects F, R, G, NB, H, SB and D simulated by the PROTHEUS numerical results. Transect F, near Samothrace,
corresponds to the C transect of G2T3. Transect R is along the Thermaikos Gulf, with remarkable river outflows. Transect G contains the sill between the Sporades plateau and the
Limnos shelf, connecting the NAT and the Central Aegean Basin. Transect NB is from the Lesbos shelf to the northern Cyclades plateau (NB is the northern part of the B transect). Its
southern part in the Cretan Sea is the SB transect, which goes from the Cyclades Plateau to the Eastern Cretan Sea. Transect H contain the sill between Central Chios Basin and the
Cretan Sea. Transect D, with the southern shelf of Ikaria and intermediate (~1000 m) and deep (~2000 m) points, south of the Eastern Cretan Sea.
March 1988. Numerical simulations by Wu et al. (2000); Demirov and
Pinardi (2002) among others, pointed out major winter events in the
years 1981, 1983, 1987, and 1990.
We here focus on path and formation of dense water currents during
the 1987 cold winter and spring, in order to infer recursive phenomena
of geological interest. We pair field data (i.e., hydrographic data from
G2T3) with theoretical models (i.e., the streamtube model by Smith,
1975; Killworth, 1977; Rydberg, 1980, and other viscous models by
Shaw and Csanady, 1983; Shapiro and Hill, 1997; Killworth, 2001), numerical outputs (from the PROTHEUS numerical model; Artale et al.,
2010), and satellite measurements (AVHRR time series; Josey, 2003;
Marullo et al., 1999a, 1999b) to infer dense water formation sites.
Such an approach aims to determine the possible pathway and spreading of the resulting bottom currents from the Northern Aegean Trough
and the Limnos-Lesbos Plateau till the Cyclades Plateau, also in relations
with contourite deposits in this plateau. All this is complemented by the
analysis of a large set of high-resolution seismic-reflection data from the
Cyclades Plateau and Myrtoon Basin in order to map late Quaternary
sedimentary structures. The combination of marine geological data
with oceanographic analyses provides relation between bottomcurrent activity and contourite drifts, as suggested by Rebesco et al.
(2014).
This work is structured as follows: the study area is described in
Section 2; data, methodologies, and both numerical and theoretical
models are described in Section 3. Analyses and results from numerical
simulations of temperature (T) and salinity (S) vertical transects as well
as the analysis of the theoretical model for bottom current pathways are
in Section 4, for both North Aegean Trough (NAT) and Central Basin. The
bottom sediments of the Cyclades Plateau and the Cretan Sea hydrology
are analyzed in Section 5. A short synthesis and conclusions are in
Section 6.
2. Study area
The Aegean basin has a very complex structure of different local environments as canyons, river deltas, coastal or offshore sediments
(Maley and Johnson, 1971; Zervakis et al., 2000).
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M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
In particular, the NAT consists of the North Sporades (≈ 1460 m
deep) and Athos basins (≈ 1150 m deep). At its southern boundary,
the Northern Aegean Sea is connected by the Limnos-Sporades channel
(≈500 m deep at the sill) with the Central Aegean Sea. The North Aegean Sea hydrology is characterized by a dense bottom layer, an upper
layer of Levantine Intermediate Waters (LIW) and a surface layer.
Often, a much colder and fresher surface Black Sea Water (BSW) crosses
the Dardanelles and works as a thermal insulator over the NAT (Plakhin,
1972; Poulos et al., 1997; Zervakis et al., 2000; Gertman et al., 2006).
The North Aegean Sea is characterized by a sea floor with deep
trenches, shelves and sills. About its sedimentary patterns, Hamann
et al. (2008) focused attention on the sediment origins: they analyzed
the different roles of west Anatolian fluvial suspension load, Aeolian
dust, the transport of the suspension load to the core position by the anticlockwise surface currents. Karageorgis et al. (2005) investigate, in
particular, the effect of Thermaikos Gulf rivers and currents. They find
that in that Gulf dense water flows enter from SE and then are trapped
in a cyclonic gyre, just off the shelf. Consequently, fluvial sediments
discharged from the rivers remain mostly present within the plateau.
The Central Aegean Sea (in the following Central Basin) is divided in
the northern Cyclades Plateau and the Skyros, North Ikaria and Chios basins (Fig. 1). Its hydrology is similar to that of the NAT, but with a different inflow of BSW. Indeed, modified BSW enters the Central Basin
through its NW flank and then follows the cyclonic flow of the Aegean
Sea.
Lykousis (2001) analyzed the subaqueous bed-forms of the Cyclades
Plateau and found dunes, sand waves, 2-D mega-ripples, sand ribbons.
He also remarked how such bottom structures would imply large bottom water velocities (from ≈40 to 200 cm/s) that, however, were not
recorded during his cruise, where the measured water velocity was
about 6–10 cm/s.
The South Aegean Sea is composed by the Ikaria basin, the southern
shelf of the Cyclades Plateau and the Cretan basin, the largest depression
of this region reaching ≈2000 m in depth. The southernmost part of the
Cretan Sea communicates with the Eastern Mediterranean through the
Cretan Arc straits.
3. Methods and data
3.1. In situ hydrographic and seismic reflection data
The few available hydrologic data (i.e. 31 CTD casts) during year
1987 were collected in Aegean Sea during the LIA-5-87 cruise onboard
of the R/V Aegeo. The North Aegean Sea was sampled from 27 February
to 3 March 1987 after a week of strong cold northerlies (G2T3). The hydrographic casts (Fig. 2) are analyzed in order to characterize the typical
hydrological structures and, in particular, to validate the hydrographic
information we derive from the numerical model, in terms of temperature, salinity, and potential density.
The mapping of late Quaternary sedimentary facies on the Cyclades
Plateau (Fig. 3) is based on the interpretation of extensive 10 in. airgun
seismic-reflection profiles collected during years 1986–1992 through
several oceanographic cruises with the R/V Aegaeo (Tripsanas et al.,
2015).
3.2. Satellite Sea Surface Temperature (SST) data
We analyze an Aegean subset of SST data from the AVHRRPathfinder, focusing on those regions characterized by dense water formation. These SST data have a daily resolution of ≈4 km (two times per
day) and are derived from the 4-km Global Area Coverage (http://www.
nodc.noaa.gov).
In particular, we use monthly averaged SST time series from 1981 to
2006 to infer dense water formation processes from SST anomalies. We
focus on 1987 (Fig. 4), which is considered as the first year of a cold
Fig. 2. Temperature (A), salinity (B) data in the C transect of G2T3.
M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
123
Fig. 3. Shaded bathymetric map of the southwest Aegean Sea, showing (A) the distribution of the bottom-current related seismic facies, and (B, C) examples of moats and sediment drifts in
the southwest Aegean Sea.
period (Zervakis et al., 2000) even though it is not the coldest year of the
1982–2006 time-series.
3.3. The PROTHEUS model
The numerical outputs we used in this study are from the PROTHEUS
system (kindly provided by Dr. Sannino), i.e., a coupled model composed by the RegCM3 atmospheric regional model and a MITgcm
ocean model (Artale et al., 2010). The coupling of these two models is
accomplished by the OASIS3 coupler (Valcke and Redler, 2006; for
more details see Appendices A and B).
We stress that, in being a hydrostatic model the PROTHEUS is not
particularly suitable for simulating dense water chimney phenomena.
Despite this source of uncertainties, we use T and S vertical transects
from this model to analyze the main hydrographic setting (and their
monthly evolution) in some crucial areas of the Aegean and Crete
Seas, an approach that is not largely affected by the non-hydrostaticity
of the model.
The PROTHEUS simulation is performed on an area including the entire Mediterranean Sea, where there are also considered inflows from
rivers and external seas, as Dardanelles that is here schematized as a
runoff. As an example, monthly averages of transect F are shown in
Fig. 5A and B, to be compared with the G2T3 measurements in Fig. 2A
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M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
Fig. 4. Monthly satellite data of winter-sping SST over the Aegean Sea during year 1987.
and B. We moreover note how the PROTHEUS bathymetry is often particularly smoothed in comparison with the real sea bottom (for more
details see Appendix B).
The model validation (Appendix B) shows that the PROTHEUS
monthly averaged temperatures are about 1–2 °C lower for the surface
layer of BSW (simulated as a layer of cold fresh water often less than
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125
Fig. 5. Temperature (A), salinity (B) data for the F transect of the PROTHEUS numerical simulation. White, vertical lines represent the G2T3 hydrological casts shown in Fig. 2.
However, along this motion the vein slightly deepens (a second order
effect of friction and entrainment) and its cross section increases. In addition, just under the main streamtube a small parallel flow of slightly
denser current often occurs.
In more detail, since the streamtube has some turbulent characteristics due to the bottom friction, its motion can erase adjacent waters, that
is, a streamtube thickness h in the point x along its path evolves as
40 m thick, Figs. 2A and 5A). This large discrepancy probably is due to a
poor parameterization of the Dardanelles outflow of BSW in the
PROTHEUS model and/or a numerical bias (Artale, personal communication). For the other layers less deep than ≈500 m this kind of discrepancy is ≈0.5 °C. On the other hand, PROTHEUS salinity is about 0.9 psu
lower for BSW and then is rather realistic, perhaps a 0.1 psu lower for
the largest depths, while the PROTHEUS density is rather realistic on
the surface and good for the deep layers.
We remind that those of PROTHEUS are Eulerian data that concern
local values in a given month, and thus the Lagrangian evolution of
the previous measured values, although fully advected by the Eulerian
model, can be affected by uncertainties related to Lagrangian chaos
(Palatella et al., 2014). We also point out that the use of the Eulerian velocities from this model, which can be potentially used for our analysis,
are not suitable for our goal since bottom numerical currents are usually
affected by large errors (due to, e.g., coarse vertical resolution, lack of salinity data for model initialization) and thus might result unrealistic
(Sannino and Artale, personal communications).
where v is the ambient water velocity (Turner, 1973) and E is the entrainment parameter, usually considered as a function of the Froude
number Fr. In the Mediterranean Sea E is ~ 10−4–10−5 (Baringer and
Price, 1997; Falcini and Salusti, 2015).
Different schematizations, due to Shaw and Csanady (1983) and
Shapiro and Hill (1997), correlate the evolution of a dense water initial
patch with the bottom density gradients, thus estimating a density current flowing horizontally with speed (Nof, 1983)
3.4. Theoretical models
uNof ¼ s g 0 =f
Models for dense water evolution over a slope describe that a dense
water, immediately after its formation, can either start a slow geostrophic motion with velocity u or it can downflow as rather quick a cascade. In this first case the pioneering streamtube model (Smith, 1975;
Killworth, 1977, among others) describes the geostrophic motion of a
vein of dense water along the slope. The balance of Coriolis force and
buoyancy in a stratified fluid gives a dense water moving approximately
along the isobaths (Smith, 1975; Killworth, 1977; Estournel et al., 2005).
where s is the sea bottom slope, gʹ is the reduced gravity and f is the
Coriolis parameter.
This kind of models also allows computing the small dense water
deepening. Killworth (2001) applied a Zilitinkevich and Mironov
(1996) analysis on meteorological and oceanographic downflows
based on a Froude number Fr characterization, and found that such outflows often follow a simple trajectory characterized by a constant increase of the current depth z, so obtaining dz/dx ≈ constant. This
u dh=dx ¼ Eju−vj
ð1Þ
ð2Þ
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M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
Froude number model gives Fr ~ 1, namely h gʹ ≈ u2 and also dz/dx b 1/
400. These results are of interest since they are apparently independent
from entrainment, detrainment and details of the bottom drag. On the
other hand such approach mostly concerns bottom currents crossing a
deep sill and thus its application to currents flowing over a slope may
be questionable.
4. Analyses and results
6, 7, 8; for details see Fig. 1 and Table 1). Results are synthesized in 3
heuristic “monthly T/S plots”, namely an overlap of T/S diagrams relative to the same geographical points but for different months. In these
plots we indicate different waters as XX YYY (e.g., F 40 or NB 100, see
Table 1), where XX is the name of the transect in Fig. 1 and YYY is the
water depth (m).
Fig. 6 shows the PROTHEUS monthly averages of T/S diagrams of
NAT during January, February and March of the year 1987. We focus
on the transects shown in Fig. 1 and Table 1:
4.1. The hydrographic data
North of the Samothrace island, the C transect of G2T3 (Fig. 2) — simulated by the F transect of PROTHEUS — is about 50 km at the east of the
W transect of G2T3. The core data of the transect C of G2T3 (Fig. 2a and
b: T ≈ 13.50 °C, S ≈ 38.75 psu, σθ ≈ 29.25 at z ≈ 170 to 350 m), and the
more westward transect W (not shown, T ≈ 12.50 °C, S ≈ 38.40–
38.75 psu, σθ ≈ 29.23 at from ≈ 150 to 380 m) reveal processes of
dense water formation over the North Samothrace shelf. G2T3 observed
the densest transect L over the more southern Limnos-Lesbos Plateau
(not shown, T ≈ 13.60–13.80 °C, S ≈ 38.95 psu, σθ ≈ 29.33).
In these G2T3 sections, the typical hydrological structure of the
Aegean Sea is revealed by temperature and salinity (Fig. 2a, b),
with relatively high values of T and S at intermediate and deep
layers. In the surface layer the cold and fresh BSW, often thinner
than ≈ 30–40 m, can be seen (Zervakis and Georgopoulos, 2002;
Pazi, 2008).
During the same period, Zervakis et al. (2000) synthesized the available marine water densities near Limnos island as σθ ≈ 29.32, while further south in the Cyclades, at ≈660 m depth, is σθ ≈ 29.18; at east of the
Island of Crete was found σθ ≈ 29.16 at 1570 m depth and σθ ≈ 29.15;
west of that island, at ≈1270 m depth. In general, Zervakis et al. (2004)
interpret the deep Cretan Water as a mixing of LIW and dense water of
NAT origin, a characterization in agreement with our results.
i) the Samothrace shelf transect (F);
ii) the shelf and the deepest part of Thermaikos gulf (R);
iii) the Limnos-Lesbos shelf (G).
During January 1987 the densest and coldest waters are found in the
Samothrace shelf (F 40, with a PROTHEUS abundant σθ ≈ 29.65), where
both S and σθ decrease considerably during February and March, likely
due to a BSW interaction (Fig. 6). This salinity decrease is a general effect
since it also occurs in the other superficial (G 40) and (R 40) points,
which become also remarkably cold. Either a meteorological interaction
or an interaction with BSW are likely the main causes of such a dense
water formation.
A strong cooling also occurs in the Limnos shelf (G 40), where the
densest water were measured by G2T3. PROTHEUS data in January
show T ≈ 14.2 °C and in March T ≈ 12.3 °C for the (G 40) water, a
cooling likely influenced by particularly dramatic storms during March
1987 (Lagouvardos et al., 1998).
Temperature in the deepest point of the Samothrace shelf (i.e., F
470), as well as for the Thermaikos gulf (i.e., R 760, R 360), remains
rather constant while salinity increases from January to March (Fig. 6).
Therefore, there is no evidence of any mixing between this deep (F
470) water with the overlying shelf (F 40), although some mixing of
4.2. SST patterns from remote sensing
In his careful analysis of the Aegean Sea SST, Josey (2003) recognized
the essential importance of the thermal effects for the density evolution
of the surface waters during the EMT winters, in comparison with evaporation. From January to May of 1987, the coast near Samothrace, the
Thermaikos gulf, and the Limnos-Lesbos shelf appear to be particularly
affected by strong winds from east, funneled by the Greek and Turkish
orography (Sayın and Beşiktepe, 2010 and Sayin et al., 2011). These
winds determine lower temperatures in these regions in comparison
to Central and Southern Aegean Sea. In turn, during June, the Cyclades
Plateau has rather lower temperatures (Fig. 4) than the northern area.
In addition, Fig. 4 shows that the 1987 was not particularly cold,
while the Turkish coast seems colder than the other shelves (Zervakis
and Georgopoulos, 2002; Sayın and Beşiktepe, 2010 and Sayin et al.,
2011). No particular thermal effect is evident over the Cyclades during
1987 while, in the following years, the water temperature decreases
by 1–1.5 °C (see Fig. S1 in Supplementary). This highlights the role of
the winter 1986–87 in preconditioning the EMT, then fully developed
during the early '90s (Sayin et al., 2010 and 2011).
Summarizing, dense water “sources” are recognized in the Samothrace and Limnos-Lesbos shelves, as suggested by G2T3, or in the
Cyclades Plateau as well (Zervakis et al., 2000 and 2002; Sayin et al.,
2010 and 2011; among others). Although in the Thermaikos Gulf
surface dense waters are evident, this gulf is also affected by a
remarkable fresh water input from rivers (Estournel et al., 2005).
4.3. Numerical results in the NAT basin [transects F, R, and G]
In order to simulate lacking field data and to assess dense vein pathways and dynamics, we analyzed monthly averaged T and S data in
some selected transects as obtained from the PROTHEUS model (Figs.
Fig. 6. Monthly T/S diagrams for the “critical points” in NAT during January, February and
March. These diagrams are the overlap of T/S diagrams during January (segment
beginning in the T/S diagram), February (intermediate point of the segment) and March
(segment with an arrow point) for the year 1987. The March surface point (F 40), that
corresponds to the G2T3 measurements, is out of this diagram because of its surprising
values T ≈ 9.4 °C, S ≈ 37.5 psu and σθ ≈ 29.9, where some complex interaction of the
local water with the BSW looks probable. Colors indicate the corresponding transect in
Fig. 1.
M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
127
northward to balance the massive southward flow of (G 40) and (G
200) waters, could explain the deep (F 470) time evolution.
In turn, the points R in the Thermaikos Gulf have a rather irregular
behavior, probably due to March storms or local river outflows, while
only (R 300) looks mainly mixed with (F 40). These mixings support
the G2T3 idea that the (F 40) dense water moved westward along the
isobaths over the shelf break in a slow geostrophic motion. Of particular
interest is that all densities in the southern sector of the basin are smaller than those of (F 40) and (G 40) during January, namely σθ ≈ 29.65
and 29.43 (Fig. 6).
4.4. Numerical results in the Central basin [transects G, NB, H, and D]
We now analyze the following PROTHEUS monthly averages of T/S
diagrams of the Central basin during January, February and March of
the year 1987 (Fig. 7) in:
Fig. 7. As Fig. 6, monthly T/S diagram of transects of the Central Basin during January
(segment beginning in the T/S diagram), February (intermediate point of the segment)
and March (segment arrow) for the year 1987. Colors indicate the corresponding
transect in Fig. 1.
(F 470) is possible with (G 40), i.e. the L transect of G2T3, or with the sill
water (G 200).
About the Limnos-Sporades channel, the point (G 200) “moves”
towards — appears to be mixed with — the denser (G 40) water. Indeed
(G 200) increases its potential density of ≈ 0.1 in February ≈ 0.2 in
March, finally reaching σθ ≈ 29.42 in April. This agrees with the hypothesis of a robust interaction between (G 40) and (G 200) waters (Fig. 6).
In the next we will show that also the (F 40) water has a similar interaction with (G 200), but with some months of delay. On the other
hand, a mixing of (F 470) with the (G 200) sill water, partially flowing
Fig. 8. As Fig. 6, monthly T/S diagram of the Central Basin transects during April (segment
beginning in the T/S diagram), May (intermediate point of the segment) and June
(segment arrow) for the year 1987. Colors indicate the corresponding transect in Fig. 1.
i) the (G 200) and (G 40), already discussed above;
ii) the southern boundary of the Central Basin, i.e. transect H with
its eastern point at 40 m depth (H 40), near Bodrum (Fig. 1)
and its sill (H 250);
iii) the point (D 200) near the island of Ikaria, East of the Cycladic
shelf;
iv) the NB transect (part of transect B from north of the Cyclades till
Lesbos), i.e. points (NB 50) East of Lesbos, (NB 100) near the island of Andros, at the northern Cycladic border, and the deep
sill (NB 500).
Fig. 7 shows that in these months the near-surface points (G 40) and
(H 40) become remarkably colder and fresher, mainly in March, likely
due to the strong front over the Aegean Sea (Lagouvardos et al., 1998).
The deep (H 250), (NB 500), and (D 200) have a common evolution,
i.e. a rather steady T and an increasing S from January to March, indicating a mixing with the salty, deep current flowing northward along the
Turkish coast (Gertman et al., 2006). The increase of salinity might be
induced by upwelling due to strong winds blowing off the Turkish
coast (Zervakis and Georgopoulos, 2002; Sayın and Beşiktepe, 2010
and Sayin et al., 2011). However, the lack of a corresponding cooling
in the eventual upwelling region (Crépon et al., 1984) suggests that a
mixing with salty waters is the most important effect (Figs. 7 and 8).
In this period the (NB 50) and (NB 100) waters move as — and are
mixed with — the cold and dense (G 200) and (G 40) waters arriving
from the NAT. Based on the estimate of mixing rates, we argue that
the relative dynamics shown in Fig. 7 mark the presence of both northern cold waters in this NB transect. The (G 40) point with PROTHEUS
σθ ≈ 29.43 and (G 200) with σθ ≈ 29.37 have the largest water
densities.
In synthesis, during winter a large amount of northern cold water arrives in the Central Basin from (G 40) and the (G 200) sill, influencing
the (NB 50) and (NB 100) points (Fig. 9 and Table 2). To balance such
southward flow the (G 200) water shows a counter-flow through the
Limnos-Sporades sill towards the deep (F 470) basin, as evidenced by
the (F 470) evolution in Fig. 6.
During spring (i.e., April, May and June), a strong warming of the superficial (G 40) and (H 40) waters is evident, till reaching ≈ 18 °C in
June (Fig. 8). An unexpected, thin surface layer of fresh water with
S ≈ 38.61 psu is present over the point (NB 100), near the Cyclades
shelf during May and June, perhaps of the BSW origin. All other points
have mild variations: the superficial (NB 50) water shows a modest
temperature increase while the sill (H 250 m) water becomes slightly
fresher. Finally we notice that the (D 200 m) water is colder, with a
S ≈ const while, during winter, S had a strong increase.
Regarding dense waters near the Cycladic Plateau, in the Central
Basin we observe that (Fig. 9 and Table 2): the (NB 100) water evolves
from σθ ≈ 29.37 (March) to ≈ 29.19 (June) while (NB 50) goes from
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Table 1
Characteristics of the PROTHEUS transects. In red two points now shown in Fig. 1 (black dots).
Transect
Orientation
Origin point
End point
Bottom slope
Critical points and depths (m)
Point symbol
F
N–S
25°11ʹE 40°58ʹN
25°11ʹE 40°06ʹN
Length (Km)
94
5 × 10−3
R
NW–SE
23°06ʹE 39°56ʹN
24°14ʹE 39°27ʹN
115
170
7 × 10−3
G
E–W
25°33ʹE 39°60ʹN
24°13ʹE 39°21ʹN
8 × 10−3
NB
NE–SW
25°42ʹE 39°26ʹN
24°53ʹE 38°00ʹN
165
83
232
H
E–W
27°34ʹE 37°11ʹN
25°16ʹE 37°11ʹN
204
2 × 10−3
D
N–S
26°05ʹE 37°13ʹN
27°04ʹE 35°01ʹN
340
4 × 10−3
SB
NE–SW
24°25ʹE 37°08ʹN
22°53ʹE 35°05ʹN
357
6 × 10−3
Northern shelf, 40
Bottom, 470
Northern shelf, 40
Intermediate, 300
Bottom, 760
Sporades, 360
Limnos, 30
sill, 200
Lesbos shelf, 50
deep sill, 500
Cyclades shelf, 100
Turkish shelf, 20
deep Ikaria, 250
Cyclades plateau,100
South Ikaria, 200
Cyclades, 1000
deep Cretan, 2000
Southern Cyclades, 800
western Cretan,1500
F 40
F 470
R 40
R 300
R 760
R 360
G 40
G 200
NB 50
NB 500
NB 100
H 20
H 250
H 100
D 200
D 1000
D 2000
SB 800
SB 1500
5 × 10−3
Fig. 9. Pathways of dense water masses in the context of the streamtube (see text). It also indicated the position of points (NB 50), (NB 100) and (NB 500). Green arrows represent the
evolution of the (F 40) water; yellow arrows represent the evolution of the dense (G 40)–(G 200) water; red arrows represent the northward evolution of the (H 250), (NB 500), and (D
200) deep waters.
M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
129
Table 2
Estimated temperature (T), salinity (S) and potential density (σθ) values in NB at different depth (50, 100 and 500 m) in SI.
Months
January
February
March
April
May
June
T
σθ
S
NB50
NB100
NB500
NB50
NB100
NB500
NB50
NB100
NB500
14.51
13.79
13.20
13.80
13.70
14.28
14.53
14.20
13.02
13.49
13.35
13.01
13.21
13.26
13.40
13.51
13.36
12.98
38.88
38.85
38.89
38.78
38.81
38.82
38.92
38.93
38.80
38.83
38.62
38.61
38.71
38.75
38.84
38.84
38.83
38.84
29.07
29.21
29.37
29.14
29.21
29.08
29.10
29.17
29.37
29.19
29.14
29.19
29.21
29.25
29.24
29.20
29.34
29.36
σθ ≈ 29.37 to ≈29.08 in June during the same time interval. By considering that the deep (NB 500) water at the sill had a density ≈29.24 during January–March, we can argue that this water probably mixes with
(G 200) and reaches σθ ≈ 29.34 in May; σθ increases to ≈ 29.36 in
June, very similar to the σθ ≈ 29.4 observed by G2T3.
The possible candidates for the origin of such density changes are
those waters coming from the Limnos-Sporades sill (G 200) and shelf
(G 40), or a deep northward current along the Turkish coast of salty
(H 250) water (Fig. 1), or just cold storms. In this regard, the (H 250)
sill water evolution shows that its salinity indeed increases from
~ 38.72 psu in January to ~ 38.86 psu in April and May, and then in
June it decreases to S ~ 38.78 psu, somehow similar to the other southern (D 200) deep point. This implies a decrease of the deep sea salty
northward transient in June. An effect of this (H 250) water in the Cyclades (NB 100) water mass looks difficult, since this water is warmer
and saltier than the (NB 100) water.
The remarkable density increase during May–June of (NB 500) is
likely due to the overlapping of all these cold flows: first from (G 200)
water, mostly in May, then for the summer an interaction in June–July
with dense waters of (G 40) origin, while no direct effect of storms on
the (NB 500) water is particularly evident. In synthesis, deep (NB
500 m) water in the Cyclades plateau appears to be mixed mainly
with cold water from NAT and Limnos-Lesbos Plateau, as suggested in
the early studies of Zervakis et al. (2000, Zervakis and Georgopoulos,
2002).
4.5. Numerical results in the Cretan basin [transects SB and D]
About the onset of the main EMT outflow from the Cretan Sea,
PROTHEUS T/S data (transects SB and D) show that the point (SB 800)
has a slightly varying T ≈ 14.4–14.6 °C and S ≈ 38.82–38–38.86 psu
and the σθ ≈ 29.04–29.08. For (D 2000; in the Cretan Basin, Table 1)
the salinity is around 38.73 psu, T ≈ 13.55 °C and the density is
≈ 29.06–29.08. For the western (SB 1500), the temperature is
T ≈ 13.74 °C while S ≈ 38.74–38.76 psu and the density ≈29.17 always
with very small variations. This suggests that deep layers south of the
Cyclades Plateau and in the deep Cretan Sea both salinity and temperature show little variations, and are similar to field data quoted in
Zervakis et al. (2000). From these critical points we infer that no outflow
of the Aegean dense water is evident during the January–June 1987
period.
water mixes with the (G 200) in the region around the LimnosSporades channel and, in the following months, will flow towards
south-west (Fig. 9), following the general cyclonic current of the Central
Basin (G2T3; Zervakis et al., 2000).
By considering the distance from the shelf where (F 40) to the
Limnos-Sporades sill (i.e., ≈400 km) and a water velocity of ≈4 cm/s,
as estimated by G2T3, it would take ≈ 120 days (Figs. 1 and 6) to
cover such a distance. Similar velocities can be estimated over the
Samothrace shelf from applying the Nof velocity in Eq. (1). Along this
hypothetic pathway (Figs. 6 and 9) the mean bottom slope in between
the Athos and Sporades sector (i.e., a distance of ≈ 100 km) would
give a delay of about 90 days, implying a March–April dense water arrival in the Limnos-Sporades channel. Velocities estimated by using the
Killworth model (Eq. (1)), which are not affected by the local bottom
characteristics, are much larger and reach values of ≈ 30 cm/s. This
would imply an earlier arrival of dense (F 40) water at the LimnosSporades channel during February (Table 3).
In addition, the application of the streamtube model (Eq. (2)) to the
Samothrace transect C of G2T3 (i.e., transect F of the PROTHEUS transects) gives a dense water westward velocity u ≈ 4–5 cm/s. Considering
the small velocities of these dense water streamtubes in the NAT, their
thickness h is expected to increase as dh/dx ≈ E/2 ≈ −10−5. We remark
that these time estimates, however, are affected by significant uncertainties due to the presence of several local, topographic effects that
cannot be easily considered here.
From Eq. (1), we also obtain that the thickness h of the (F 40) water
layer increases as dh/dx ≈ E/2 along its ≈400 km long pathway in the
NAT shelf-break. By assuming a constant E/2, we obtain that entrainment gives an increase of h around 15 m at the end of its travel. This
small variation is neglected in the following discussion, although some
are considerable and not-captured mixing processes (e.g., confluence
of different water masses towards a similar pathway) is certainly present. Finally, from the streamtube model we estimate a deepening of this
(F 40) cascading of ≈1 m/km.
Regarding the Central basin, T/S diagrams suggested a northward, high salinity flow along the Turkish coast. The (G 200)
water is about 230 km (NB 500; Fig. 1) and, by considering a shelf
slope of ≈ 8 × 10 − 3 , a local u Nof ≈ 8 cm/s (see also Lykousis,
2001) would give one month of delay from the (G 200) density
peak in April (Figs. 7, 8, and Table 3), giving a first cold water arrival
in (NB 500) May. From the Limnos shelf, at ≈ 480 km from the (NB
4.6. Velocities and pathways from the theoretical models
About the marine water dynamics in the NAT, we notice that the (G
40) densest water (that formed south of Limnos) moves geostrophically
along the isobaths and, through the time, reaches similar T-S values of
the (G 200) in-channel water (Figs. 1 and 6). For the (F 40) water, the
data agree with the streamtube dynamics: a near geostrophic pathway
along the NAT shelf-break (Figs. 6 and 9). This can explain the observed
lack of mixing between (F 40) and (F 470). This westward motion along
the isobaths receives also some support from the dynamic topography
observed by Gertman et al. (2006) in the northernmost part of the
NAT during 1988. Therefore, it is reasonable to assume that the (F 40)
Table 3
Estimated values of the streamtube model velocities along transects in SI.
Transect
gʹ
s
h
Nof velocity
Killworth
velocity
F
R
G
NB
H
D
SB
10−3
8 × 10−4
10−3
10−3
5 × 10−4
5 × 10−4
10−3
5 × 10−3
7 × 10−3
8 × 10−3
5 × 10−3
2 × 10−3
4 × 10−3
6 × 10−3
80
60
70
100
70
50
50
0.08
0.06
0.08
0.05
0.01
0.02
0.06
0.3
0.2
0.2
0.3
0.2
0.2
0.2
130
M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
Fig. 10. PROTHEUS monthly data of density in the NB vertical transect, during the January–June 1987 period.
500) point with a uNof ≈ 5 cm/s, one has about 3 months of delay
and thus a dense water arrival can be estimated in June–July (Fig.
10).
4.7. Geophysical results
The upper 200 m of the sedimentary packet in the Cyclades Plateau
most probably is the product of late Quaternary sedimentary processes
(Anastasakis et al., 2006). It consists of the six seismic facies shown in
Fig. 3. Parallel, continuous, mid to low-amplitude seismic reflections
represent hemipelagic sediments. In turn high to mid-amplitude, discontinuous to incoherent reflections and overlying on a hard substrate
with no further sound penetration, are usually interpreted as shallow
marine and fluvial sediments. Lenses of incoherent mid- to lowamplitude reflections interbedded in hemipelagic sediment represent
mass-transport deposits (MTDs). Seismic facies related to bottomcurrent are expressed by erosional channel-like features, which are
bounded on their sides by mounded seismic packets, consisting of
high-amplitude incoherent to semi-parallel, discontinuous mid- to
low-amplitude seismic reflections. Similar seismic facies have been described from many ocean environments around the world (Rebesco
et al., 2014, and references therein) and are usually interpreted as
moats and sediment drifts. These sedimentary structures mark the
main entrance of the deep water we investigate here between the
Islands of Euboea and Andros (Channel B in Fig. 3). Moreover, they
occur along some interesting pathways towards the Myrtoon Basin
and the Cretan Sea.
5. The Cyclades plateau deposits
We now relate the above results on dense water hydrodynamics and
pathways with the evidence of bottom sediments in the Cyclades Plateau (Fig. 3). According to the distribution of the sediment drifts, the
bottom currents that affect their formation processes enter the Cyclades
Plateau through a strait in the North-East sector of this basin
(i.e., between the Islands of Euboea and Andros, NB 100 point in Figs.
1 and 9) and exit towards west and southwest through three deep
straits in the Myrtoon Basin and the Cretan Sea. This flow is depicted
by the sedimentary structures (sand waves, dunes, and sand ribbons)
already identified by Lykousis (2001) in the same area. The continuation
of the sediment drifts at the much larger water depth of the Myrtoon
Basin indicates that the bottom currents exhibited a strong cascading
feature, suggesting that they have large densities.
We notice that bottom velocities, as obtained from the streamtube
model, are rather small (Table 2) with respect to those that are needed
to form such sedimentary structures (Rebesco et al., 2014). However,
this region is characterized by a very complicated bottom topography
that have a remarkable influence on the bottom water velocity and
their temporal evolution. It could indeed happen that impulsive strong
currents, affected by irregular bottom confinements, give origin to
strong nonlinear phenomena. Indeed, Lykousis (2001) from the analysis
of such sedimentary structures predicted velocities up to 1–2 m/s while
his own measurements were 7–10 cm/s, in some agreement with our
estimated velocities (Table 3). Moreover, streamtube models deal
with steady motions, while the Lykousis (2001) bottom structures
probably captured only the paroxismical stage of the currents, which
might not have lasted for a long time.
Finally, we stress that in the streamtube case lower velocities are expected because we focus our analysis just on a single year, prior to the
EMT event. Moreover, Lykousis (2001) estimates would be, in turn,
rather similar to the Killworth model velocities. Therefore, these sedimentary results somehow support the model here proposed which, accordingly, agrees with the southward dense water outflow off the
Cyclades Plateau. There, one can indeed remark at the boundary of the
Cyclades Plateau and the Myrtoon basin (Fig. 3) some evidence of bottom current dynamics, a particular, interesting feature that suggests future investigations.
6. Conclusions
In this paper, we analyze the thermodynamics of Aegean Sea currents during the year 1987, just before the onset of the EMT event
(Roether et al., 2013).
M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
The complexity of the EMT event has been largely analyzed by several authors. G2T3 found very dense water masses and assumed that
the densest waters of the Cretan Sea arrived via slope convection from
the Cycladic Plateau shelf, although specific dense water paths were
not identified.
By combining satellite, numerical, field, and theoretical information, we here complement and what was found in seminal work
(e.g., Zervakis et al., 2000; Theocharis et al., 1999a, 1999b; Lascaratos
et al., 1999; Malanotte-Rizzoli et al., 1999; Nittis et al., 2003) by presenting a novel viewpoint about the Aegean Sea dense water dynamics. Our
analysis considers the northern dense water formation zones observed
by Zervakis et al. (2000) in the light of a stream tube model (Smith,
1975; Killworth, 1977).
From SST satellite data, we recognized the main sources of very
dense water in the Samothrace shelf and the Limnos-Lesbos Plateau.
From T and S data, simulated by the PROTHEUS numerical model, we
found that the densest water from the Samothrace shelf flowed
geostrophycally along the NAT isobaths, as a streamtube, towards the
Limnos-Sporades channel. There, it mixed with a dense water from
the Limnos-Lesbos Plateau and moved geostrophically towards the Cyclades during April, May and June with different pathways and arrival
times. In addition, it is reasonable to argue that the heavy storm that
occurred on March 1987 had a significant role on this dynamics. A different high-salinity, northward flow along the Turkish coast, probably
related to wind induced upwelling (Sayin, 2010 and 2011), seemed to
play a secondary role. Our stream tube approach essentially supports
the Zervakis et al. (2000) analysis.
Our findings were then related to the presence of sedimentary structures in the area. Results suggested that the main entrance of dense
water in the Cyclades Plateau is between the Islands of Euboea and Andros. Moats and sediment drifts in this area agree with this southward,
dense water outflow. This opens new questions about the downflow towards the Myrtoon Basin and the deep Cretan Sea. Indeed, regarding the
EMT temporal evolution, there was no evidence of dense water outflow
off the Cretan Sea towards the Levantine basin during the January–June
1987 period. Since such a dense water outflow was observed at the end
of the 1987, we can argue that the main EMT event started after the
period of our analysis.
Supplementary data to this article can be found online at http://dx.
doi.org/10.1016/j.margeo.2016.01.012.
Acknowledgments
We are deeply grateful to Dr. Volfango Rupolo, who started this
study, to Dr. Artale, Gacic, Mosetti for creative criticism. We are also
indebted to Dr. Marullo for the satellite data, help and criticism and to
Dr. Sannino for providing the PROTHEUS data, while Dr. Garra and
Gasbarra were of great help in an early period of this work. We also
thank both the anonymous reviewers for their support and their constructive criticisms. This work has been supported by the Italian National Flagship Project RITMARE (Marine Italian Research), funded by the
Italian Ministry of Education, University and Research within the National Research Program 2011–2013.
Appendix A. The PROTHEUS numerical model
In addition to the text information, here we remind that the
RegCM3 (i.e., the athmospheric component of PROTHEUS system)
is a 3-dimensional, sigma-coordinate, primitive equation, hydrostatic regional climate model (Giorgi et al., 1993a, 1993b, Giorgi and
Mearns, 1999; Pal et al., 2007). Air-sea exchanges are treated using
the parameterization of Zeng et al. (1998). The configuration used
here has a uniform horizontal grid spacing of 30 km on a Lambert
conformal projection and 18 σ-levels are used. The simulation is performed on an area including the entire Mediterranean Sea. Lateral
boundary conditions are supplied every 6-hours by interpolating
131
horizontal wind components, temperature, specific humidity and
surface pressure from a global atmospheric model.
The ocean component is based on the MITgcm developed by
Marshall et al. (1997a, 1997b); here the model is used in its hydrostatic,
implicit free-surface, partial step topography formulation (Adcroft et al.,
1997). The model has a resolution of 18 18 , equivalent to rectangular
meshes of variable resolution with the meridional side of about 14 km
and the zonal one ranging from about 9 km in the northern part of the
domain to about 12 km in southern part. The model has 42 vertical Zlevels with a resolution varying from 10 m at the surface to 300 m in
the deepest part of the basin, and an intermediate resolution of about
40–50 m between the depths 200–700 m. Horizontal viscous and
diffusive terms are modeled with a bi-harmonic formulation with
diffusivity and viscosity coefficients equal to 1.5 × 1010 m4/s. Vertical eddy-diffusivity is modeled via a Laplacian formulation with
diffusivity coefficient ranging from 3.0 × 10 − 5 m 2 /s at surface
to 1.0 × 10 − 7 m 2 /s at the bottom, while the viscous coefficient is
hold constant to 1.5 × 10 − 4 m 2 /s over the whole water column
(Sannino et al., 2009). Deep convection is simulated through the enhancement of the vertical diffusivity to about 1 m 2 /s in regions
where the stratification becomes unstable.
Respect to the version of the MITgcm model described in Artale et al.
(2010), in this version natural boundary conditions for salinity are used,
that is P + R − E (precipitation plus runoff minus evaporation) is treated as a real fresh water flux. Monthly river discharges computed from
the RegCM3 total runoff are used by the oceanic model. Catchment
basins for 147 rivers falling into the Mediterranean Sea have been reconstructed using Total Runoff Integrated (TRIP) database. Rivers
discharging in the Black Sea have been collected to dense fresh water
fluxes that reach the Mediterranean Sea via the Dardanelles Strait. A
monthly value for the discharge of Black Sea is then included as a further
river. The initial condition for the oceanic run is obtained from a previous run performed using a 3D relaxation towards the MEDATLAS-II
climatology.
Appendix B. The PROTHEUS model validation
As a check of the numerical simulation over the Aegean Sea, we
compared monthly SSTs produced by PROTHEUS with Optimally Interpolated SST (OISST, Marullo et al., 2007) from Pathfinder 5.0 SST data.
The agreement between the two time series (Fig. S2) confirms that
PROTHEUS is able to reproduce reasonably well the annual SST cycles,
though some minor discrepancies between the model and the data do
not exceed 1.5 °C and less than 1 °C in average. The difference between
OISST and field data shows no time drift (Marullo et al., 2007), as also
happens between the model and OISST data and we therefore deduce
that the model has no remarkable drift in the overlapping period.
We moreover compare the PROTHEUS hydrologic transect (F 40)
with CTD data for the North Samothrace transect C (Fig. 2) of G2T3 during March 1987. There is a low surface temperature T ≈ 11.20–11.50 °C
(in PROTHEUS simulation T ≈ 9.5 °C), most probably due to a thin surface layer of BSW ≈ 40 m thick. Just below one has T ≈ 12.5 °C while in
the deepest layers T reach 13.5 °C (PROTHEUS has T ≈ 12.6 °C down to
100 m, remaining rather homogeneous with 12.8 °C at the sea bottom).
As remarked by a referee it is possible that the BSW from Dardanelles is
underestimated till a factor 2 by the PROTHEUS model.
About other parts of the Aegean in the same period, Zervakis et al.
(2000) synthesize the available field data at Limnos (near Samothrace)
with T ≈ 13.30 °C (T ≈ 12.8 °C in the PROTHEUS data) while more
South in the Cyclades at 660 m depth T ≈ 14.25 °C (T ≈ 13.4 °C in the
PROTHEUS data), at east of the Island of Crete T ≈ 14.05 °C at 1570 m
depth (T ≈ 13.6 °C in the PROTHEUS data) and T ≈ 14.10 °C West of
that Island, at just 1270 m depth (T ≈ 13.8 °C in the PROTHEUS data).
In synthesis the PROTHEUS simulation is about 2 °C low for the BSW,
rather poorly simulated, and ≈0.5 °C low for the other layers.
132
M. Bellacicco et al. / Marine Geology 375 (2016) 120–133
In the same F transect, salinity S ≈ 38.00 psu in the BSW surface
layer, rapidly increasing with depth to 38.25–38.50 psu till reaching
≈38.80 psu in the deepest layers (PROTHEUS simulates the BSW layer
as low as S ≈ 37.0 psu on the surface, then S ≈ 38.5 psu at 60–100 m
depth, to a rather homogeneous value S ≈ 38.75 psu below 200 m).
About other parts of the Aegean in the same period, Zervakis et al.
(2000) synthesize the available field salinity at Limnos (near Samothrace) with S ≈ 38.84 psu (S ≈ 38.82 psu in the PROTHEUS data)
while more South in the Cyclades at 660 m depth S ≈ 38.94 psu
(S ≈ 38.95 psu in the PROTHEUS data), at east of the Island of Crete
S ≈ 38.84 psu at 1570 m depth (S ≈ 38.8 psu in the PROTHEUS data)
and S ≈ 38.85 psu West of that Island, at just 1270 m depth
(S ≈ 38.78 psu in the PROTHEUS data). Again, there is a PROTHEUS salinity about 1 psu low for BSW and then rather realistic values, perhaps a
0.1 psu low for the largest depths.
The surface BSW density is less than ≈29.05 while the PROTHEUS
density is ≈29,0 in the BSW layer.
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