Dense water formation in the Aegean Sea: Numerical simulations during the Eastern Mediterranean Transient

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. C9, 8120, doi: /2002jc001352, 2003 Dense water formation in the Aegean Sea: Numerical simulations during the Eastern Mediterranean Transient Kostas Nittis National Centre for Marine Research, Institute of Oceanography, Athens, Greece Alex Lascaratos Department of Applied Physics, University of Athens, Athens, Greece Alex Theocharis National Centre for Marine Research, Institute of Oceanography, Athens, Greece Received 19 February 2002; revised 10 April 2003; accepted 30 April 2003; published 9 September [1] Dense water formation processes in the Aegean Sea (eastern Mediterranean) are studied using a three-dimensional numerical ocean model. The simulations cover the period during which major changes that affected the thermohaline circulation of the whole Mediterranean Sea were recorded. Sensitivity studies that focus on the role of freshwater budget are presented, and the results are evaluated against available hydrological data of the same period. The very cold winters of 1987, 1992, and 1993 and the extended dry period that affected the whole eastern Mediterranean Sea are the main driving mechanisms, corresponding to 50% and 32%, respectively, of the excessive deepwater volume formed in the Aegean after The reduced Black Sea Water outflow during the same dry period was another important forcing mechanism, contributing 18% to the total formation, while the increased inflow of saline waters from the Levantine Sea after 1992 was an additional preconditioning factor. The locations and mechanisms of water formation processes are identified with combined analysis of data from the March 1987 oceanographic cruise in the Aegean Sea and the respective model results for that period. Deep water is found to be formed mainly through open ocean convection in the central and north Aegean Sea, while the contribution of shelf areas is limited. Intermediate water is also formed through open ocean convection in the southern Aegean Sea during cold winters as well as in the central and northern Aegean during mild winters. The total volume of dense water formed during corresponds to an annual formation rate of 0.24 Sv for deep water and 0.34 Sv for intermediate water. INDEX TERMS: 4243 Oceanography: General: Marginal and semienclosed seas; 4255 Oceanography: General: Numerical modeling; 4283 Oceanography: General: Water masses; 4504 Oceanography: Physical: Air/sea interactions (0312); 4532 Oceanography: Physical: General circulation; KEYWORDS: dense water formation, eastern Mediterranean, Aegean, numerical simulations Citation: Nittis, K., A. Lascaratos, and A. Theocharis, Dense water formation in the Aegean Sea: Numerical simulations during the Eastern Mediterranean Transient, J. Geophys. Res., 108(C9), 8120, doi: /2002jc001352, Introduction [2] The thermohaline circulation of the Mediterranean Sea is driven by its negative freshwater and heat budget that transforms the relatively warm and fresh surface Atlantic water to the colder and more saline Mediterranean intermediate water. The main sites of dense water formation in the basin are the Gulf of Lions where the Western Mediterranean Deep Water is formed, the Adriatic Sea where the Eastern Mediterranean Deep Water is formed, and the NW Levantine basin where the Levantine Intermediate Water (LIW) is formed [Wüst, 1961]. The Aegean Sea Copyright 2003 by the American Geophysical Union /03/2002JC has been also proposed as source of dense water with intermediate or deep characteristics [Pollak, 1951; Miller, 1963]. Intermediate water of Aegean origin was found in isolated lenses or thin layers just below the LIW [Schlitzer et al., 1991] and was named Cretan Intermediate Water (CIW). Recent studies [Malanotte-Rizzoli et al., 1999; Roether et al., 1999] attribute a much more significant role to this warm and saline water mass, that seems to occupy a thick layer below the LIW in large areas of the eastern Mediterranean Sea. [3] Since the beginning of 1990s a large number of oceanographic surveys have monitored a significant change in the deep thermohaline circulation of the eastern Mediterranean [Roether et al., 1996]. The deepest parts of the basin have been filled by very dense and young water of PBE 21-1

2 PBE 21-2 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA Aegean origin that lifted up the older bottom waters of Adriatic origin changing in few years the water mass structure that was constant since the existence of observations. The event is known as the Eastern Mediterranean Transient (EMT). The reanalysis of hydrological data from the 1980s and 1990s in the Cretan Sea showed that the increased formation of dense water in the Aegean started during 1987 and was intensified during the following years, reaching maximum values during [Theocharis et al., 1999]. It was found to be a combined effect of salinity increase and temperature decrease in the water masses of the Aegean. The volume of the dense water that fills the deep basins of the south Aegean (Cretan Deep Water (CDW)) increased gradually, occupying also the intermediate and upper layers of the Cretan Sea by It was followed by a massive outflow toward the eastern Mediterranean where it occupied the bottom layers. [4] Analysis of data from the north Aegean also supports the hypothesis that changes were initiated during the extremely cold winter of 1987 [Zervakis et al., 2000]. The reduced Black Sea Waters (BSW) outflow in the Aegean Sea during the early 1990s is suggested to be an additional mechanism that facilitated the dense water formation process. The surface layer of BSW that occupies a large area of the north Aegean acts as an insulator of the underlaid saline waters of Levantine origin and thus prevents, in many cases, dense water formation; the weakening of this mechanism enhanced dense water formation in the area. [5] Lascaratos et al. [1999] reviewed hydrological observations of the last decade that describe the EMT and presented preliminary results of numerical studies that attempt to reproduce the processes and the mechanisms responsible for it. According to these experiments, dense water with both intermediate and deep characteristics is formed in the Aegean Sea depending upon the intensity of the atmospheric forcing. The period after 1986 is characterized by increased formation of dense water during almost every winter. According to their numerical experiments, the dense water is stored in the deep basins of the south Aegean and outflows to the eastern Mediterranean first through the eastern (deeper) straits of the Cretan Arc and then trough the western (more shallow) straits. [6] Numerical studies of Samuel et al. [1999] suggest that an intensification of northerly winds in the Aegean Sea between 1987 and 1993 increased the exchanges through straits of the Cretan Arc leading to increased inflow of saline LIW and outflow of dense water from the Aegean. In their experiments this water had intermediate characteristics and was not dense enough to reach the bottom layers of the eastern Mediterranean. A different mechanism was proposed by Wu et al. [2000], who used a 1/8 resolution model of the Mediterranean Sea to study the impact of cold winters in the north Aegean. Their experiments were idealized using a strong cooling of the north Aegean for 8 consecutive winters ( ). According to their results, the cold water formed in the north Aegean Sea mixes with the more saline waters of the Cretan Sea and produces realistic water properties at the bottom layers of the eastern Mediterranean Sea. Their budget analysis showed that the redistribution of salt from the upper 1000 m to the deep layers can explain more than 75% of the simulated salt increase at these levels. [7] Stratford and Haines [2002] used a combination of the mechanisms proposed by Samuel et al. [1999] and Wu et al. [2000] as a better approximation to the observed meteorological anomalies. They applied the artificial excess winter cooling only for years 1987, 1992, and 1993 and a stronger northerly wind component during Their results indicate that the role of wind is secondary to the role of buoyancy forcing and that the cooling of these specific 3 winters can initiate the deepwater formation process. They conclude that the roles of freshwater and heat fluxes are equally important and there is a need for better representation of freshwater budgets in future simulations. [8] Compared to previous modeling studies that used artificial perturbations to initiate the EMT and mainly focused on the resulting circulation and hydrological changes in the eastern Mediterranean, the present work focuses on the mechanisms of dense water formation inside the Aegean and their interannual variability, incorporating in the same time a more realistic representation of heat and water fluxes. It is an extension of the work of Lascaratos et al. [1999] using this time a higher-resolution grid in order to identify in detail and differentiate the mechanisms in the various sites and estimate their relative importance. Furthermore, this study presents additional sensitivity experiments in order to assess the relative importance of the different freshwater anomalies. Hydrological observations from the Aegean Sea are also used for a direct validation of model results and a more detailed description of the winter 1987 event. [9] The numerical model, the meteorological data and the air-sea interaction scheme are presented in section 2. Results from the numerical simulations of the period, together with the different sensitivity experiments are presented in section 3. A more detailed analysis of the winter 1987 data and model results is given in section 4, where the sites and mechanisms of dense water formation inside the Aegean Sea are identified. Finally, in section 5 the work is summarized and the final conclusions are presented. 2. Numerical Experiments 2.1. Model Setup [10] The Princeton Ocean Model (POM) developed by Blumberg and Mellor [1987] is used in this study. The same model has been successfully used in the Mediterranean Sea for various applications: general circulation studies [Zavatarelli and Mellor, 1995], water mass formation studies [Lascaratos and Nittis, 1998], regional modeling of the Adriatic and the Aegean Sea [Nittis et al., 1995; Zavatarelli and Pinardi, 2003]. It has been also used for a real time oceanographic nowcast/forecast system of the Mediterranean [Horton et al., 1997]. [11] The main characteristics of POM are as follows. POM is a free surface model; that is, the surface elevation is a prognostic variable. It uses bottom following sigma coordinates in the vertical. A time-splitting technique is used to calculate the 2-D and three-dimensional equations with different time steps. It includes the Mellor and Yamada [1982] 2.5-order turbulence closure scheme for the computation of vertical mixing coefficients. Finally, horizontal

3 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA PBE 21-3 Figure 1. Model grid and corresponding topography (contour interval 200 m). diffusivities are calculated according to the Smagorinsky formula. [12] The model grid covers the whole eastern Mediterranean excluding the Adriatic Sea (Figure 1). Two different implementations of the model have been used: one high resolution (grid spacing 10 km) and one low resolution (20 km). The coarse grid has been used mainly in the customization experiments where the different boundary conditions and atmospheric forcing schemes were tested. Both configurations use 19 sigma levels in the vertical with logarithmic distribution in the upper layer. Preliminary results from the low-resolution experiments have been presented by Lascaratos et al. [1999], while the results presented here are exclusively from the high-resolution implementation. [13] A side effect of the horizontal diffusivity formulation used by POM is the weak relaxation to climatology [Mellor, 1998]. In climatological studies this has a positive effect since it minimizes model s climate drift [Ezer and Mellor, 2000]. On the other hand, in studies of interannual variability this term leads to dumping of perturbations and thus weakening of the interannual signal. The effect is minimized in high-resolution applications where the Smagorinsky diffusion term becomes very low. An alternative way to reach the same result is to reduce the coefficient in the Smagorinsky formulation. This approach has been used in the present study and the Smagorinsky coefficient was set equal to 0.02, i.e., five times smaller than the typical value of 0.1. This procedure reduces but does not eliminate the dumping effect of the horizontal diffusion term. [14] At each grid point the depth was computed using bilinear interpolation of the DBDB5 bathymetric database of the U.S. Navy that has a 1/12 of a degree resolution. The pressure gradient error is a well know problem of sigma coordinate models when steep topography is present [Haney, 1991]. In POM the problem is handled by subtracting an area mean density profile before the computation of horizontal density gradients. This technique, together with smoothing of topography can eliminate the error [Mellor et al., 1994]. In this application the topography was smoothed with a third-order Shapiro filter [Shapiro, 1970] applied twice, followed by a selective filter that smoothes depth variations exceeding 20% between adjacent grid points. The model was finally tested through experiments with zero external forcing and no initial horizontal density gradients. The error introduced in both grids was found to be negligible since the residual kinetic energy was less than m 2 /s 2 after 300 days of integration Boundary Conditions [15] The open boundaries of the model are located: a) between Sardinia and Africa; b) between Sardinia and Italy (Tyrrhenian Sea); c) at the Otranto Strait. The first two describe the exchanges with the western Mediterranean and the third with the Adriatic Sea. For salinity, temperature and the tangential component of velocity, an upstream advection equation is used. When there is inflow through the open boundary, temperature and salinity is prescribed from the seasonal climatology of MODB-MED5 that is also used for model initialization [Brasseur et al., 1996]. A free radiation and a Sommerfeld [Blumberg and Mellor, 1987] radiation condition are used for the normal to the boundary depthintegrated and baroclinic velocities, respectively. Under mean climatological forcing, the total volume exchange rate estimated by the model is 0.89 Sv for the Sicily strait and 0.61 Sv for the Otranto strait with a seasonal modulation of ±0.3 and ±0.15 Sv respectively. Experimental estimates vary between 0.6 and 1.2 Sv for the Sicily strait [Grancini and Michelato, 1987] but the latest results are closer to the upper limit of this range (1.1 Sv according to Astraldi et al. [1996]). The observational estimates for the Otranto strait are close to 1 Sv [Gacic et al., 1996] but values 3 times smaller had been reported in the past [Orlic et al., 1992]. [16] The Dardanelle s strait that controls the exchange between the Black Sea and the Mediterranean is also treated as open boundary but with different conditions. A two-layer system is explicitly prescribed at this strait with inflow of fresh BSW in the upper layer and outflow of saline Aegean waters in the lower layer. A similar scheme was used by Horton et al. [1997], while previous applications [Zavatarelli and Mellor, 1995] have treated those straits as a river; that is, only the inflow was prescribed. In the present model setup, both inflow and outflow transports follow a seasonal cycle with maximum values during June, but the net transport (inflow-outflow) is kept constant to 300 km 3 /yr (900 km 3 /yr inflow minus 600 km 3 /yr outflow). The salinity of the upper

4 PBE 21-4 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA Figure 2. Annual mean precipitation fields from different sources: black bars are central and south Aegean average values from Theocharis et al. [1999]; shaded bars are Aegean Sea average values from Global Precipitation Climatology Centre (GPCC), combined product GPCP version 1c [Rudolf et al., 1994]; and open bars are eastern Mediterranean average values from GPCC. The mean value is from Tselepidaki et al. [1992]. and the lower layer is fixed to 29.5 and 38.9 respectively following Unluata et al. [1990] Parameterization of Air-Sea Fluxes [17] The model incorporates realistic air-sea interaction parameterization using an interactive estimation of buoyancy fluxes that allows representation of one-way (ocean to atmosphere) feedback [Rosati and Miyakoda, 1988; Lascaratos and Nittis, 1998]. Instead of using predefined air-sea fluxes the model computes them using the estimated SST at each time step and the given atmospheric parameters. The atmospheric data are from the hours ECMWF reanalysis. The model uses the two components of surface wind speed W u, W v, the air temperature T a and dew point temperature T DP at surface and total cloud cover C TCC. Relative humidity is computed from T a and T DP using the equation for saturation vapor pressure. The bulk formulae used in this study are the Bignami formula for long wave back radiation [Bignami et al., 1995], the Kondo formula for sensible and latent heat flux [Castellari et al., 1998], while for short wave radiation the formulas of Rosati and Miyakoda [1988] were followed. [18] The weak point of surface boundary conditions is the freshwater flux (evaporation minus precipitation, E P) that has to take into account precipitation, a parameter that is usually poorly estimated. In contrast to the other atmospheric parameters that come from a self-consistent data set (ECMWF reanalysis), precipitation data are taken from climatological database [Jaeger, 1976]. This means that no interannual variability signal is introduced by this parameter. There is, nevertheless, evidence of strong variability in precipitation all over Greece during the study period. In fact, the years marked the driest period in Greece during this century [Tselepidaki et al., 1992], while equally dry were the years of [Theocharis et al., 1999]. In order to better quantify the reduced precipitation for these years, the rain gauge data of Theocharis et al. [1999] were reanalyzed selecting averages of 4 stations of the central and south Aegean. For these stations multiyear mean values ( ) are available from Tselepidaki et al. [1992]. The results (Figure 2) indicate that the reduction in rainfall during the years and is approximately 40% compared to the multiyear average (300 mm/yr compared to 500 mm/yr). In order to include this significant anomaly in the model s forcing, a reduction of 40% was introduced in the climatological fields of precipitation for the years and This anomaly was used to modify the monthly climatological values in the Aegean Sea and was equally applied to all the months of the respective years. [19] Precipitation data from the Global Precipitation Climatology Centre (GPCP) [Rudolf et al., 1994] were also analyzed, and annual mean values averaged over the Aegean and the eastern Mediterranean Seas were estimated for the years The results indicate that the extended dry period has not affected only the Aegean Sea but the whole eastern Mediterranean as well (Figure 2). This variability could not be properly quantified, since no records are available for longer periods and thus a comparison against multiannual means cannot be made (as it was done for the Aegean Sea where the means are available). Therefore no adjustment was introduced on the climatological precipitation fields of the eastern Mediterranean. Finally, in order to incorporate the interannual variability of BSW outflow in the Aegean Sea suggested by Zervakis et al. [2000], a 25% reduction was introduced on transport through Dardanelle s for years and a 15% for years [20] The results presented in the next section are based on three numerical experiments that simulate the 16year period The central experiment (EX1) uses the abovedescribed configuration. In the other two experiments, the parameterizations of freshwater anomalies have been gradually removed in order to study the relative impact of each forcing mechanism. Experiment EX2 is without the anomaly of BSW outflow, while in EX3 the precipitation anomaly has also been removed. The differences between these experiments start after 1989 when these anomalies are first introduced. The model is initialized by the results of a 20 year long climatological experiment that used a perpetual year forcing computed from the monthly average (climatology A described by Lascaratos et al. [1999]). With this procedure, a repeating cycle has been

5 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA PBE 21-5 Figure 3. Annual mean atmospheric forcing parameters, estimated air-sea fluxes, and volumeintegrated salinity in the Aegean Sea, as reproduced by the 16 year numerical experiments. The total buoyancy flux is estimated for the north Aegean Sea, for comparison with Zervakis et al. [2000]. Thick lines correspond to EX1, thin lines respond to EX2 (without the anomaly of BSW outflow), and dotted lines respond to EX3 (with neither BSW nor precipitation anomalies). achieved before the application of the realistic 6 hours forcing for the period. 3. Variability During the Years Surface Buoyancy Fluxes [21] Annual mean values of air-sea fluxes calculated from the 3 numerical experiments are presented in Figure 3. These time series clearly depict the years of 1987 and as the periods of maximum heat loss ( 25 to 30 W/m 2 instead of a mean value of about 10 W/m 2 ) with secondary maxima during 1981 and These periods are attributed to the cold winters of 1987 and and the periods of maximum wind speed during the early 1990s. The increased wind speed has a clear impact on the evaporation as well: the two years with maximum wind speed (1987 and 1994) give the maximum E P budget in the EX3 experiment where no precipitation anomaly has been applied. [22] The precipitation anomalies introduced in EX1 and EX2 have a clear impact on the E P time series since they represent a more than 30% increase of this budget for the respective years and Their effect is very important on the volume integrate salinity of the Aegean Sea, while it is rather minor on the temperature time series (not presented in Figure 3). Without any freshwater anomaly (EX3) the salinity time series presents local maxima during the heat loss periods ( , , and ) with a reducing trend in between. This trend is reversed when the freshwater anomalies are introduced (EX1 and EX2) and increasing volume-integrated salinity values are observed after The relative increase of salinity between EX3-EX2 and EX2-EX1 indicates that the effect of the imposed BSW reduced inflow is approximately 30% less important than the reduced precipitation mechanism, at least for the volume-integrated salinity values in the Aegean Sea. [23] Although the model uses the best meteorological fields available for that period and an advanced scheme for heat and water flux estimates, we cannot exclude a possible underestimation of buoyancy losses especially during extreme events. It is known that wind speed is underestimated by the ECMWF analysis [Komen et al., 1994] especially in the Mediterranean where local orographic effects are smoothed out by the relatively low-resolution meteorological model grid. Furthermore, the interactive scheme used for the estimation of buoyancy fluxes describes only the oneway feedback since it uses a variable SST but a fixed air temperature provided by the ECMWF reanalysis. This analysis has taken SST variations into account but for larger space scales. The best way to estimate such errors is through direct comparison to meteorological data from

6 PBE 21-6 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA buoy stations that do not include the usual bias of landbased stations. Such a comparison was carried out for the Aegean Sea during the period when buoy data were available from 10 stations [Nittis et al., 2001] and showed a 5 23% underestimation of wind speed even if a higher-resolution meteorological model is used. A relatively good agreement was found for air temperature estimates (bias 1 8 %). [24] Even if similar reference data were available for the study period, a large uncertainty still remains because of the use of bulk formulas. Therefore the only way to assess our flux computations is through comparison to other independent estimates. The first comparison is against long-term means. The selection of bulk formulas was based on their ability to give results comparable to the long-term means for the Mediterranean Sea [Castellari et al., 1998] and the specific application area; the mean value of heat loss for the model domain (eastern Mediterranean excluding the Adriatic) was estimated to be 11.7 W/m 2 and the mean E P budget 590 mm/yr in good agreement (although the E P is close to the lower limit) with previous estimates based on completely different data sets [Garrett et al., 1993; Bethoux and Gentili, 1994]. [25] The second assessment is made using the annual mean net buoyancy flux defined as: B ¼ c 1 w gaq T þ gbsðe PÞ where c w the specific heat of water, g the gravity acceleration, a the thermal expansion coefficient of seawater, b the salinity contraction coefficient and S the salinity [Gill, 1982]. The above term was computed for the north Aegean Sea (north of 38.5 N) and was compared to similar estimates of Zervakis et al. [2000] on the basis of the COADS data set. The BSW outflow (as it is expressed in our model configuration) was transformed to precipitation equivalent and was added to the precipitation term of the equation. The results (Figure 3) follow very close those derived by the COADS data in both absolute values and temporal variability. The most important differences are depicted during years for which the estimated values are higher than those derived by Zervakis et al. [2000] especially for the central experiment that includes the freshwater anomalies. The same applies for 1991 and 1994 while for years our estimates are slightly lower than those derived from COADS data. For years , when the total heat loss is relatively low, the freshwater anomalies are found to play the leading role on the total annual buoyancy loss of the north Aegean Formation Sites and Rates [26] The dense water formation process is often visualized using the depth of selected isopycnal surfaces. When an isopycnal that characterizes intermediate or deep layers outcrops at the surface then the corresponding layer is ventilated and we have formation of dense water. For the visualization of the water formation events reproduced by the 16 years numerical experiments, the depth of the isopycnal surface kg/m 3 during the period of maximum average surface s q was selected (Figure 4). In most cases this period is between end-february and mid-march but during winter 1987 maximum densities were reached during the last 2 weeks of March while another strong convection event took place in January of that year (see also section 4). [27] The size of the outcropping area inside the Aegean Sea has significant variations, with minimum extent during the very mild winters of 1984 and 1986 and maximum values during the very cold years of the late 1980s and early 1990s, when it covers the whole Aegean and NW Levantine. The central Aegean (the Cyclades plateau and the Skyros-Chios basins) and the shelf areas along the NE coasts of the basin are found to be sites where ventilation of intermediate layers is taking place during almost every winter (see Figure 9b for a more detailed map where these areas are marked). The western Cretan Sea is also an area where deep mixing is reproduced by the model in many occasions after It should be noted that using the isopycnal kg/m 3, the formation of both intermediate and deep water is taken into account. Indeed, the CIW is usually characterized by s q kg/m 3, while the s q value of kg/m 3 has been considered as the lower limit of the Cretan Deep Water [Theocharis et al., 1999]. [28] In order to quantify the above-described interannual variability of dense water formation in the different subregions of the Aegean Sea, the annual formation rates were estimated for the south (35 37 N), central (37 39 N) and north (39 41 N) Aegean (Figure 5a). These estimates use the results of the central experiment (EX1) following the methodology of Lascaratos and Nittis [1998] for two different s q classes: for s q > kg/m 3 (deep water) and for s q in the range kg/m 3 (intermediate water). The total dense water formed during is 3.8 Sv yr for deep water and 5.39 Sv yr for intermediate water. The central Aegean Sea is the primary formation site for deep water (56% of the total volume) while the south Aegean is the primary site (59%) of intermediate water formation. The deepwater formation rates range between 0.1 and 0.4 Sv in the central Aegean but they exceed 0.5 Sv during the extreme winters of when maximum formation rates are also observed in the north and south basins. In fact, during 1993 the annual deepwater formation rate exceeds 1.3 Sv for the whole Aegean Sea with important contribution from all 3 subregions. Intermediate water is formed regularly in the central Aegean with a typical formation rate of Sv while in the south Aegean the formation is less regular but more intense, with maximum rates exceeding 0.5 Sv during the winters of 1983, 1987, , and [29] The deepwater (s q > kg/m 3 ) formation rate was also estimated separately for the shallow areas of the Aegean Sea (depth < 100 m) in order to quantify the contribution of these sites in the whole process (Figure 5a). The results indicate that the contribution of these areas is indeed very limited since the estimated volume of deep water during is 0.15 Sv yr, i.e., less than 5% of the total volume formed in the Aegean Sea. This indicates that the open sea convection is the primary process of dense water formation in the Aegean Sea. [30] For a better assessment of dense water formation mechanisms during the period , annual mean hydrological properties in the 3 subbasins were estimated for the layers m, m, and 600 m-bottom (Figure 5b). These estimates show the contrast between the

7 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA PBE 21-7 Figure 4. Maps of the depth (m) of isopycnal kg/m 3 during the period of maximum average surface s q (in windows of 10 days), for each year of simulation EX1. Black color marks the areas where dense water is formed (outcropping of the isopycnal).

8 PBE 21-8 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA Figure 5. (a) Time series of dense water annual formation rates (Sv) and (b) hydrological properties in the south (35 37 N, thick solid line), central (37 39 N, thin solid line) and north (39 41 N, dotted line) Aegean Sea. The formation rates have been estimated for deep (s q > kg/m 3 ) and intermediate (s q kg/m 3 ) waters, as well as for shallow areas (depths < 100 m). The hydrological properties are volume-integrated annual means and have been estimated for three different layers: m, m, and 600 m-bottom. south warm and saline basin with the north cold and fresh Aegean. The central Aegean is a transitional zone between these two extremes and its properties are, in most cases, closer to those of the southern basin. Maximum densities at intermediate and deep layers are observed in the central and northern parts, where s q exceeds and kg/m 3 respectively by the end of The interannual variability of properties verifies that the excessive dense water formation was a combination of temperature decrease after 1986 and salinity increase after It is interesting to note that the temperature decrease is prominent in all 3 layers and subbasins, especially in the south and central Aegean. On the other hand, the salinity increase is very pronounced at the northern and central basins but practically absent from the upper and intermediate layers of the southern Aegean. This contrast is attributed to a) the fact that BSW anomaly influences primarily the north Aegean and b) the important exchanges through the straits of the Cretan Arc that make the upper layers of the southern Aegean an integrated part of the eastern Mediterranean Sea; the precipitation anomaly signal which is introduced only over the Aegean sea is easily removed from the upper layers of the southern basin, while the temperature anomaly that influenced a large area of the eastern Mediterranean is sustained. [31] The salinity increase in the central basin leads to higher values than the south basin at intermediate and upper layers during the period This is an important preconditioning mechanism for dense water formation and it explains the increased formation rates in this area in combination, of course, with the increased wintertime buoyancy loss. The rate of salinity increase is even stronger in the north Aegean, throughout the water column. In fact, by the end of the simulation period the intermediate and deep salinity values in the north basin are very close to the values of the central and south Aegean which indicates an important contribution from this basin during the last phase of EMT. [32] According to the results presented in Figures 4 and 5a, increased water formation rates are reproduced not only during the cold winters of 1987 and but also during the period in between which was characterized by relatively mild winters. This is in agreement with observations of Theocharis et al. [1999] and should be attributed to the preconditioning of winter 1987 and the increased buoyancy loss due to the freshwater anomaly. Indeed, the strong formation event of 1987 had a strong impact on the vertical structure of the Aegean and facilitated dense water formation in the following years. The evolution of s q structure in the Cretan Sea (Figure 6) shows a much weaker stratification in January 1988 compared to the same month of In 1988 there is clear presence of new dense water in the deep layers and the isopycnal kg/m 3 that hardly appears in the deep parts of the western basin is now at an average depth of 800 m. As it is already suggested by Figure 4, the western part of the south Aegean is an important source of dense water and in this area the kg/m 3 isopycnal is already at a depth of 600 m in By 1991 the volume of dense water in the deep layers of the Cretan sea have considerably increased and the isopycnal kg/m 3 is above the sills, at a depth 500 m, which means that dense water already outflows toward the eastern Mediterranean. The outflow in the previous years has intermediate water s q characteristics kg/m 3. In 1994 the maximum s q has increased to more than kg/m 3 and the outflow at sill depth is above kg/m 3. [33] Apart of the preconditioning role of winter 1987 the increase of buoyancy loss due to freshwater anomalies of was the main mechanism of intense water formation during that period of relatively mild winters. In fact, earlier modeling studies of EMT based exclusively on heat flux anomalies did not reproduce deep convection in the period between 1987 and 1991 [Stratford and Haines, 2002]. In 1989 the water anomalies give a 50% increase on the buoyancy loss (Figure 3, comparison of EX1 and EX3) while in 1990 the buoyancy flux is increased by 3 times when water anomalies are introduced. This mechanism is also present during the second phase of freshwater anomaly ( ) that coincides with the period of maximum cooling.

9 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA PBE 21-9 Figure 5. (continued) [34] In order to quantify the contribution of freshwater anomalies, the deepwater (s q > kg/m 3 ) formation rates for the whole Aegean Sea were estimated for the 3 different experiments (Table 1). The total dense water production after 1987 corresponds to 3.3 Sv yr, which is approximately half of the volume estimated by Roether et al. [1996]. This volume is reduced to 2.7 Sv yr in EX2 (no reduction of BSW outflow) and to 1.6 Sv yr in EX3 (no reduction of BSW and no precipitation anomaly). This indicates that the freshwater anomalies account for almost half of the total volume of dense water formed in the Aegean Sea between 1987 and 1994 and the impact of precipitation anomaly is approximately 2 times larger than the impact of reduced BSW outflow (1.1 compared to 0.6 Sv yr, respectively) Role of Horizontal Advection [35] An important question on the EMT forcing mechanisms is the role of advective fluxes through the straits of the Cretan Arc. Malanotte-Rizzoli et al. [1999] suggest that changes in general circulation structures in the eastern Mediterranean during the beginning of 1990s were responsible for a change of the LIW spreading pathway. The westward propagation of LIW was blocked by the intensified anticyclonic structures south of Crete (Ierapetra and Mersa-Matruh Gyres). It was suggested that this blocking led to increased inflow of saline LIW in the Aegean Sea through the eastern straits of the Cretan arc, which was a strong precondition for dense water formation during the following years. [36] Our numerical simulations successfully reproduce most of the anticyclonic circulation features in the southern part of the eastern Mediterranean, namely the Mersa-Matruh Gyre southeast of Crete, the Shikmona Gyre southeast of Cyprus and the anticyclonic gyre in the Ionian, as well as their intensification between 1987 and Nevertheless, the Ierapetra Gyre, which is the strongest anticyclonic structure of the Mediterranean Sea circulation and is usually attached to the SE coast of Crete, is not realistically reproduced by the model because of the luck of highresolution atmospheric forcing [Horton at al., 1997]. This means that the suggested blocking mechanism is not reproduced by our experiments. Indeed the time series of transports through the straits of the Cretan Arc (Figure 7) do not show any increasing trend in the period (in fact there is a weak negative trend in the eastern straits during ). The simulated total transports through the eastern straits are close to 2 Sv of inflow and 1 Sv of outflow resulting in a net inflow of 1 Sv. In the western straits we have the opposite picture with similar mean values and a net outflow of 1 Sv. This is in agreement with the general circulation pattern of the Aegean Sea that leads to a large-scale cyclonic circulation with net inflow from the eastern straits and net outflow from the western. The above transport values are also in agreement with similar estimates from in situ observations [Kontoyiannis et al., 1999]. [37] The most interesting feature in the transports time series in the increase of inflow and outflow at both straits after On the basis of the comparison with the results of Figures 3 and 5a, this can be attributed primarily to the increased outflow rates of dense water through the deep layers of the straits and the compensating increase of inflow at the upper layers and secondarily to the role of wind stress. In fact, all local maxima during (which are

10 PBE NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA Figure 6. West-east s q (kg/m 3 ) cross sections in the central Cretan Sea for January 1985, 1988, 1991, and 1994 showing the evolution of density structure in the basin. more pronounced in the western straits) are attributed to the dense water formation events of 1983, , and , with the exception of the 1981 maximum that can only be correlated with the intensified wind patterns of that year. As it is already noted, the increased inflow has a positive feedback on dense water formation process but according to the above results this mechanism has played an important role during the last phase of EMT (after 1992). Nevertheless, since the model is not able to reproduce the blocking effect of the early 1990s, the role of a similar mechanism during the preconditioning phase of the formation event cannot be excluded. In fact, the poor representation of that mechanism might be partially responsible for the underestimation of deepwater formation rates described in the previous paragraph Validation of Model Results [38] The availability of repeating oceanographic observations in the Aegean Sea during the study period allows a qualitative and quantitative evaluation of numerical model results. The first validation is based on estimates of volumeintegrated hydrological properties in the deep layers of the Cretan Sea. The deepwater properties of this basin were selected because a) similar estimates based on in situ observations are available from recent publications and b) this area is a reservoir where dense water, formed in various sites of the Aegean Sea, is stored before its exit through the Cretan Arc straits. When the volume of this dense water is increased and its interface is raised above the sill depth then it outflows in the eastern Mediterranean Sea. This comparison gives an overall view of the model s efficiency since deep waters properties are affected by Table 1. Annual Formation Rate of Aegean Deep Water (s q > kg/m 3 ) for EX1 (Central Experiment), EX2 (Without the Anomaly of BSW Outflow), and EX3 (With Neither BSW nor Precipitation Anomalies) Annual Formation Rate, Sv Year EX1 EX2 EX

11 NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA PBE Figure 7. Time series of annual mean transports (Sv) through the straits of the Cretan Arc: inflowing (solid thick line), outflowing (solid thin line), and net inflow (dotted). model results and observations (apart of the period). This difference appears right from the beginning of the time series and should be attributed to the initialization fields. The slight increasing trend of temperature between 1987 and 1992 is also reproduced but underestimated by the model. [40] The increase of deepwater salinity occurs in short jumps that follow the major dense water formation events (1983, 1987, 1989, 1990, 1992, and 1993). This means that during the months that follow these events the deep layers of the basin are filled with new saline water. With the exception of the newly formed water that renews the deep layers of the Cretan Sea is also slightly warmer. The fast rise of temperature and salinity is followed by a slower decrease until the next event of dense water formation occurs (even if it is the next winter). This is probably related to diffusive or advective mixing of the newly formed dense water with the ambient water masses. This variability can be also detected in the time series of observations when data from successive cruises with short time difference are available. [41] The evolution of deep temperature and salinity properties for the other two experiments (EX2 and EX3) verifies that the freshwater anomalies have indeed played an important role. When they are omitted, both temperature and salinity variability is strongly underestimated compared to the central experiment. Furthermore, it is confirmed that the correction applied on the precipitation is more drastic than the one applied on the BSW outflow. In experiment EX3, the small increasing trend of temperature between various processes (convection, spreading, mixing) and the memory of these processes is sustained for a long period. A more qualitative validation is carried in the next section where model results of winter 1987 are compared against in situ hydrological data collected by a basin-wide oceanographic survey in the Aegean Sea during the same period. [39] For the calculation of CDW characteristics, properties were integrated between 1000 m and the bottom every 1 month of the period. The resulting time series are presented in Figure 8 together with results of the same calculation applied on hydrological data from the Cretan Sea during the period [from Theocharis et al., 1999]. The model results (central experiment EX1) follow the evolution of both temperature and salinity observations, but tend to underestimate the two major variations that describe the evolution of dense water formation during these years: the continuous increase of salinity after 1987 and the big temperature drop after The observations indicate an overall increase of salinity by 0.2 between 1987 and 1995, while the model results give an increase of 0.11 for the same period and 0.16 when the upper limit is taken from 1993 when the model reproduces the maximum salinity. The observed temperature drop between 1992 and 1995 is 0.32 C while the simulated is 0.18 C. An almost constant difference of approximately 0.1 C exists between Figure 8. Time series of volume-integrated temperature ( C) and salinity (psu) at the deep layers (below 1000 m) of the Cretan Sea from the numerical experiments: EX1, continuous line; EX2, dotted line; and EX3, dashed line. The results of the same calculations, but using in situ hydrological data from oceanographic cruises in the area during , are marked by open circles [following Theocharis et al., 1999].

12 PBE NITTIS ET AL.: DENSE WATER FORMATION IN THE AEGEAN SEA Figure 9. Maps of surface (5 m) s q (kg/m 3 ) derived from (a) CTD data of the winter 1987 LIA cruise, (b) model results (EX1) for 1 March 1987, and (c) model results for 15 March and 1992 is reversed to a decreasing trend, while the salinity increase is underestimated for almost Winter 1987 Formation Event [42] A more detailed analysis of the winter 1987 model results in conjunction with hydrological data collected during the same period in the Aegean Sea are presented in this section. CTD data were collected in February, March, and April 1987 during a cruise of R/V Aegaeo in the framework of the international program POEM and the national program LIA [Theocharis and Georgopoulos, 1993]. The winter of 1987 can be characterized as the coldest winter of the 1980 s and it is well know for intense formation of dense water in the Aegean Sea [Zervakis et al., 2000] and the Levantine basin [Nittis and Lascaratos, 1999]. In the same time, it is the only period for which basin-wide winter observations are available for the Aegean Sea during the EMT period. [43] Surface (5 m) s q fields calculated from objective analysis mapping of the CTD data together with model estimates for 1 and 15 March 1987, are presented in Figure 9. A direct comparison between in situ data and model results must take into account that the data were collected during a wide time frame of more than two months. The northern Aegean Sea was sampled during end February and beginning of March, the central and southern part during middle to end of March, the Ionian Sea during the first half of April, the Cretan passage (between Crete and Africa) during the end of April and the eastern Cretan straits during the beginning of May. Therefore the two model snapshots presented in Figures 9b and 9c correspond to the period of sampling exclusively inside the Aegean Sea. [44] The data confirm the model results (section 3.2) that two main areas of dense water formation in the Aegean Sea are the Cyclades plateau and the Skyros-Chios basins. These are the areas of maximum surface s q that exceeds 29.3 kg/m 3 by mid March These areas are also marked in Figure 4 by the outcropping of kg/m 3 during the years of dense water formation. Both of them are areas of cyclonic circulation that favors open sea convection [Marshall and Schott, 1999]. On the large scale, the general circulation of the Aegean Sea is cyclonic with inflow of Levantine water through the eastern straits, northward flow along the Asia Minor coasts and southward flow along the western coasts of the Aegean [Ovchinnikov, 1966]. An important part of the northward flow recirculates at the Skyros basin while the rest penetrates further toward the north Aegean. This recirculation creates an almost permanent cyclonic circulation in the Skyros-Chios basins. A cyclonic circulation also dominates the whole Cyclades plateau according to wintertime observations of 1987 and 1998 [Theocharis et al., 1998]. In both areas the surface salinity values are increased because of the northward transport of the Levantine Surface Waters by the above-described circulation pattern. Therefore the same dense water formation mechanism seems to be present in both areas: the presence of a permanent cyclonic gyre and the high-salinity values create the appropriate preconditioning for dense water formation during wintertime strong buoyancy loss events. The same mechanism is responsible for the formation of LIW in Rhodes cyclonic Gyre [Nittis and Lascaratos, 1998]. [45] The NE part of the basin, north of the Lemnos Island is dominated during the beginning of March by the presence of the fresh and thus buoyant BSW plume. According to the model results, two weeks later this plume has been restricted just in front of the Dardanelle s strait because of the strong buoyancy loss and mixing that increased substantially the surface density in the whole Aegean Sea between 1st and 15th of March. Indeed, during that period the average air temperature over the Aegean Sea dropped below 5 C and the total heat loss averaged over the basin reached 600 W/m 2. This exceptional event related to a passage of a cold front over the Aegean Sea [Lagouvardos et al., 1998] had a