An eddy resolving numerical study of the general circulation and deep-water formation in the Adriatic Sea

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Deep-Sea Research I 51 (2004) 921–952

An eddy resolving numerical study of the general circulation and deep-water formation in the Adriatic Sea A. Mantziafou*, A. Lascaratos Department of Applied Physics, University of Athens, University Campus-Builds, PHYS-V, Athens, Greece Received 10 June 2002; received in revised form 13 February 2004; accepted 5 March 2004

Abstract General circulation and deep-water formation (DWF) processes in the Adriatic basin in a climatological year were numerically simulated in a high-resolution (1/20th of a degree) implementation of the Princeton Ocean Model (POM). The ‘‘perpetual’’ year atmospheric data were computed from the ECMWF Reanalysis data (1
 1
) covering the period 1979–1994. The model reproduces the main basin features of the general circulation, water mass distribution and their seasonal variability. The Adriatic Deep Water exiting through the Otranto Strait is produced with two different mechanisms inside the basin: (a) by open ocean deep convection over the Southern Adriatic Pit and Middle Adriatic Pit (b) on the continental shelf of the Northern and Middle Adriatic. The estimated contributions of both mechanisms suggest that 82% of the Adriatic Deep Water is formed inside the Southern Adriatic Pit, while all the higher density water in this water mass comes from the northern regions. The role of mesoscale eddies at the periphery of the densewater chimney in the Southern Adriatic Pit was examined and their contribution to the lateral buoyancy flux, during the convection process, found to be small. The DWF rate at Otranto Strait is 0.28 Sv with sy over 29.15. The sensitivity of the DWF processes to interannual variability of the buoyancy forcing and river runoff was assessed with a number of process-study numerical experiments. In these experiments the effect of an imposed ‘‘extreme’’ buoyancy forcing during 1 year, on the DWF rates, was to modify them during the specific year, but the effects were still present in the following normal climatological year. This shows that the DWF rates and their mass characteristics at a specific year depend not only on the atmospheric conditions prevailing that specific year but on the previous year’s as well, thus leading to the concept of a ‘‘memory’’ of the basin. r 2004 Elsevier Ltd. All rights reserved. Keywords: Adriatic Sea; Numerical modeling; Climatology; Deep-water formation; Baroclinic instability

1. Introduction

*Corresponding author. Tel.: +3-0210-7276839; fax: +30210-7295282. E-mail address: [email protected] (A. Mantziafou).

The Adriatic Sea, located at the northernmost part of the central Mediterranean, is the area where, climatologically, the deep waters of the Eastern Mediterranean (EMDW) are formed (Pollak, 1951; Ovchinnikov et al., 1985; Roether

0967-0637/$ - see front matter r 2004 Elsevier Ltd. All rights reserved. doi:10.1016/j.dsr.2004.03.006

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et al., 1993; Lascaratos, 1993). It is a semi-enclosed elongated basin (B800 km long and B200 km wide), which communicates with the Mediterranean Sea and particularly with the Ionian Sea through the Strait of Otranto (75 km wide and up to 800 m deep). The dense waters, after their formation, exit through the Otranto Strait and sink to the bottom of the Ionian basin. They can be traced in the deep layers of the whole Eastern Mediterranean up to its easternmost part, the Levantine basin. Traditionally the Adriatic Sea can be separated from north to south into three sub-basins (Fig. 1). The Northern Adriatic (NA) includes the very shallow parts of the basin in the north and extends southward to the 100 m isobath. The Middle Adriatic (MA) in the central part of the basin includes the Middle Adriatic Pit (maximum depth B250 m) and extends southwards to the Pelagosa Sill. The Southern Adriatic (SA) extends from the Pelagosa Sill to the Otranto strait. The bathymetry of the latter sub-basin reaches 1200 m at the bottom of a wide depression at its center, known as the Southern Adriatic Pit (SAP).

Calculations performed with observational data show that the Adriatic Sea has a negative heat budget (mean annual heat loss to the atmosphere of B20 W/m2), while at the same time it is a dilution basin with an annual fresh water gain of 1.14 m (Artegiani et al., 1997a). These two factors have opposite buoyancy effects. The positive fresh water budget is equivalent to a heat gain of B15 W/m2, which still leaves the basin with a buoyancy loss equivalent to a heat loss of 5 W/m2. This explains in general why dense waters are formed inside the Adriatic. Another important factor is that evaporation (E) almost equals precipitation (P) and the 1.14 m/yr surplus of fresh water mainly comes from river runoff (R) (Artegiani et al., 1997a). This fresh water input is not equally distributed over the basin. It forms a coastal jet along the Italian coast, leaving the rest of the basin and particularly the area of the SAP very much exposed to buoyancy losses during winter. Recent observations (Poulain, 2001) and studies of historical data sets (Orlic et al., 1992, Artegiani et al., 1997b) as well as modeling studies of the Adriatic (Horton et al., 1997; Zavatarelli et al.,

Fig. 1. The bathymetry of the Adriatic basin in meters (c.i. 50 m).

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2002) indicate that the general circulation of the basin is mainly cyclonic. This cyclonic circulation is primarily thermohaline, driven by a combination of surface buoyancy fluxes, river discharges, and exchanges through the Strait of Otranto (Poulain and Cushman-Roisin, 2002). The impact of the wind on the general circulation can be dominant only during intense transient episodes (Malanotte-Rizzoli and Bergamasco, 1983). Such violent wind outbreaks during winter induce extended heat losses and evaporation that drive deep-water formation (hereafter DWF) events. According to available observational data a highly dense-water mass, with sy ranging from 29.4 to 29.9, is formed in the shallow NA (NADW) during winter (Zore-Armanda, 1963; Malanotte-Rizzoli, 1977; Artegiani et al., 1989). Because of the shallow bathymetry of this subbasin, the vertical temperature and salinity profiles are almost uniform during winter, when the vertical mixing dominates over the horizontal baroclinic circulation. This water mass flows along the isobaths of the western coast forming a vein (Bigniami et al., 1990a,b; Artegiani and Salusti, 1987) which splits into two branches, one descending into the MA Pit, and the other flowing over the Pelagosa Sill (Zoccoloti and Salusti, 1987). The MA bottom is filled with NADW mixed with Mediterranean water inflows along the eastern coast according to Artegiani et al. (1989), or with locally formed MA Deep Water (MADW) during winter periods of weak inflow of surface Mediterranean water through the Otranto Strait (ZoreArmanda,1963). During the period of the strongest winter cooling (15 February–15 March) the SA Deep Water (SADW) is formed over SAP in the center of the SA gyre (Zore-Armanda, 1963; Ovchinnikov et al., 1985; Orlic et al., 1992). Deep-water formed in the NA reaches the SAP concentrated in a vein, which follows the isobaths of the western coast. As soon as this vein reaches the vicinity of Bari (Fig. 1) it encounters an offshore oriented canyon, where it destabilizes, sinks and at the same time, mixes with the surrounding waters. (Bigniami et al., 1990a,b; Orlic et al., 1992). ZoreArmanda (1963) remarks that NADW and SADW occasionally mix in the SAP, while Ovchinnikov

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et al. (1985) finds NADW underlying SADW in the SAP. Zoccoloti and Salusti (1987) suggest that SADW is the result of the mixing between the small volume of NADW entering the SAP with the Modified Levantine Intermediate Water (MLIW) inflowing from the Otranto Strait. Historical data of dissolved oxygen in the bottom layer of the SA (Buljan and ZoreArmanda, 1976) indicate that deep water is ventilated every year. However cruise data in the late 1990s have shown that bottom-reaching convection in the SAP occurs rarely (Gacic et al., 2002). This implies that the deepest layers of the SAP are renewed by the much denser water originating from northern sub-basins as confirmed by the findings of Manca et al. (2002). Adriatic Deep Water (ADW) exits through the bottom layer of the Otranto Strait (500–800 m). It amounts to 30% of the total outflow and shows a major presence in winter and spring (Manca and Giorgetti, 1998). Roether and Schlitzer (1991), using chlorofluoromethane and tritium data as tracers of the newly formed deep water, computed an annual rate of 0.2970.09 Sv of dense water exiting the Adriatic through the Otranto Strait. This estimate is in agreement with the estimates of Lascaratos (1993). During the late 1980s–early 1990s an important shift occurred in the Eastern Mediterranean, whereby the Aegean Sea replaced the Adriatic as the main source of deep waters. This major event, called the ‘Eastern Mediterranean Transient’, has been attributed to important meteorological anomalies in the area as well as to changes in the circulation patterns (Roether et al., 1996; Lascaratos et al., 1999; Klein et al., 1999). The event has lasted for a number of years and the system seems not to be in equilibrium yet, although there are signs that the transient is fading away (Theocharis et al., 1999). Furthermore recent data indicate that the Adriatic is probably on its way back to its previous dominating role as the main source of deep waters for the Eastern Mediterranean (Klein et al., 2000; Manca et al., 2002). To date, model studies on the DWF processes in the Adriatic basin are very few and most of them focus on the Northern and Middle part of the basin. In particular, Hendershott and Rizzoli

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(1976), using a vertically integrated model, showed the important role of topography upon the spreading of the dense-water mass formed on the continental shelf of the NA. Pinardi et al. (1996), using POM model with an orthogonal curvilinear grid (variable horizontal resolution: from 3 to 5 km in the NA to 10–12 km in the SA) and climatological atmospheric forcing from May (1982) adequately simulated NADW formation and spreading. Vested et al. (1998), using a nonhydrostatic model and two different atmospheric data sets, those of May (1982) and ECMWF (December 1993–July 1994), investigated the rates of formation and spreading of dense water in the NA and MA. Bergamasco et al. (1999), using POM model with 3.5 km horizontal resolution in several process studies, examined the conditions under which convective mixing extends to the bottom of the MA Pit as well as the role of the rim current along the Italian coast versus the other circulation components in distributing the NADW within the basin. They claim that the stratification characteristics of the water column within MA and in particular the density gradient between the MLIW at the intermediate layer and the NADW at the bottom layer is crucial in determining the convection depth in the MA. A number of Mediterranean scale climatological model studies (Zavatarelli and Mellor, 1995; Haines and Wu, 1995; Horton et al., 1997; Myers and Haines, 2000) include the Adriatic basin, but they have coarser horizontal resolution and the DWF processes in the Adriatic are not their main focus. The present paper focuses on the issue of DWF in the Adriatic Sea and the processes involved. For that, an eddy resolving implementation of a 3D 1 numerical model (POM) with full physics at 20 th of a degree (B5 km) horizontal resolution is used. DWF processes in the Adriatic are studied through perpetual (climatological) year simulations. The contribution of the deep-waters formed in the northern parts of the Adriatic basin is estimated but special attention is given to the open ocean convection occurring in the SA, where the largest volumes of Adriatic deep waters are being formed. Furthermore the role and effects of mesoscale eddies on the DWF processes, in the

same area, is studied and the model results are compared with theoretical predictions. Finally, the sensitivity of the DWF rates to interannual variability of forcing factors is examined, through a number of process-study numerical experiments and the concept of a ‘memory’ of the ocean is developed.

2. Model set-up The present study was performed with the threedimensional(3-d), primitive-equation, free surface, sigma coordinate Princeton Ocean Model (POM), designed by Blumberg and Mellor (1987). POM contains a time splitting technique in order to solve the fully 3-d and the depth integrated equations with different time steps. It also uses the Mellor–Yamada 2.5 turbulence closure scheme (Mellor and Yamada, 1982) which provides the vertical diffusion coefficients, whereas the horizontal diffusion coefficients are computed from the Smagorinski formula (Smagorinsky, 1963). The area covered by the model grid is 37
–46
N and 12
–22
E and is resolved by 160  200 grid 1 points with a horizontal resolution of 20 th of a degree (B5 km). The model domain extends 3
south of Otranto Strait into the Ionian Sea, so as to minimize open boundary effects. The first internal Rossby radius (R) for the Adriatic Sea ranges between 4.23 km in winter and 8.79 km in summer (Grilli and Pinardi, 1998). Our calculations, using model outputs, give similar results. Mesoscale instabilities that develop during the DWF processes have a typical wavelength of L ¼ 2pR (Madec et al., 1991) which takes a value of B26.5 km in Adriatic during winter and it is resolved by the horizontal grid spacing of the model. In order to evaluate the role of mesoscale eddies in DWF, as well as the sensitivity of the basin to interannual variability of forcing factors, 1 a coarser horizontal resolution grid of 10 th of a degree (B10 km) of the same domain and a finer 1 one of 50 th of a degree (B2 km) comprising only SA, were also used. In all three horizontal grids the vertical resolution consists of 20 s-layers, logarithmically distributed near the surface. The depth of the surface layer varies from 0.08 to 5 m, and

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the bottom layer depth from 2.5 to 80 m. The value of the HORCON parameter, used in the Smagorinski formula was 0.2, 0.1 and 0.05 for the 10 km, 5 km and 2 km grid runs respectively, in agreement with values proposed by Mellor (1998). Zavatarelli et al. (2002), in their simulations of the Adriatic basin, with variable horizontal resolution ranging from 3 to 5 km in the NA, and from 10 to 12 km in the SA, used HORCON equal to 0.1. 2.1. Bathymetry, initialization and open boundary conditions The model bathymetry was taken from the US Navy Digital Bathymetric Data Base (DBDB5) using bilinear interpolation. In order to eliminate the pressure gradient error associated to s-layers over with steep topography (Mellor et al., 1994), a first-order Shapiro (Shapiro, 1970) filter was applied in the whole domain. In addition, the criterion dx H=Ho0:3 that limits the depth difference of adjacent cells was met. The minimum depth was set to 20 m. The model was initialized with summer climatological data from the Mediterranean Ocean Data Base (MODB-MED4) (Brasseur et al., 1996) which has a horizontal resolution of 14th of a degree. In Fig. 2, the initialization fields of temperature and salinity at surface (10 m) are presented. Both fields are consistent with the Adriatic climatology derived from ATOS hydrographic data set (Artegiani et al., 1997a,b). In particular, in the temperature field a front, along the basin axis, between the eastern and western basin, with warmer waters on the western and shallower part and colder waters on the eastern and deeper part of the basin, is evident. The salinity field exhibits a north–south gradient and low salinity waters dominate the NA as a result of the great river runoff of this area. At the southern open boundary, a free radiation condition for the depth integrated velocity, normal to the open boundary, and a zero gradient condition for the free-surface elevation were used in the external mode (2-d, depth integrated equations). In the internal mode (3-d, sigma layer equations) a Sommerfeld radiation condition for the internal velocity, normal to the open bound-

Fig. 2. The model initial conditions of (a) temperature in Celsious degrees (c.i. 0.2) and (b) salinity in p.s.u. (c.i. 0.1).

ary, and an upstream advection equation for temperature and salinity were used. In case of inflow through the open boundary, temperature and salinity were prescribed by the seasonal climatology of MODB-MED4. The external and internal time steps used were 9 and 900 s, respectively. The same bathymetry, initialization and boundary conditions were used for the other two grids. The only difference in the 2 km grid is that in case of inflow through the open boundaries, temperature and salinity were prescribed by 10-day means calculated by the 4th year (steady state) of the 5 km grid run. 2.2. Forcing data The surface boundary conditions used by the model are the momentum, heat and salinity fluxes. The atmospheric parameters at the sea surface (wind field, air temperature, relative humidity,

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total cloud cover), used for the calculation of the momentum and heat fluxes, were monthly averages of the 6-h European Center for Medium range Weather Forecast (ECMWF) reanalysis data for the period 1979–1994 with 1
 1
resolution. This climatological data set has already been used by Lascaratos et al. (1999). In the preparation of the wind climatology, the transition from 6 h wind data to monthly averages is not trivial. Scalar-mean values maintain the magnitude but distort the direction especially in the case of winds of opposite directions occuring within the average time period. Statistical analysis of ECMWF Reanalysis wind data (1979–1994) inside the model domain has revealed that winds of opposite directions (mainly northern and southern winds) appear with high percentage in all months. In order to have the correct magnitude of both the wind speed and wind stress and a wind stress direction close to the prevailing one, the

scalar mean of the wind speed and the vector mean value of the wind stress were estimated. The derived wind stress for February, May, August and November is presented in Fig. 3. The Bora (northeasterly wind) and Scirocco (southeasterly wind) winds, which are the dominant winds of the region, are adequately well represented in our climatology. This climatology is close to the May (1982) wind field and quite different from NMC climatology for the Adriatic basin (see Artegiani et al., 1997a). In the latter, the prevalence of the westerlies is due to the calculation method used (scalar mean), which gives the correct magnitude for the wind speed but the wrong direction for the wind velocity and stress. Another difficulty in the preparation of the wind climatology is the generally accepted problem of the underestimated ECMWF reanalysis wind data. Cavaleri and Bertotti (1997) have shown that the wind field provided by ECMWF is

Fig. 3. The wind stress climatology used, for the months February, May, August and November.

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underestimated—during stormy events—at least by 50% for the Adriatic basin. To overcome this problem Zavatarelli et al. (2002) adopted a 50% increase of the wind field in their numerical study, while Manca et al. (2002) increased it by 30% for the computation of the heat fluxes in the Southern Adriatic. After a number of sensitivity runs, in which the overall heat budget and inflow/outflow rates in Otranto Strait have been checked, we adopted an increase of the climatological wind field by 20% for this study. The bulk formulae used for the estimation of the heat budget components are presented in Table 1. The heat fluxes Qs, Qb, Qe and Qh were computed interactively from these bulk formulae, the model produced sea surface temperature (SST), and the climatological values of wind speed (scalar mean), air temperature and relative humidity at the sea surface (2 m), and total cloud cover. A number of formulae were tested and the set providing a total heat budget Q ¼ 24 W/m2, which is in agreement with calculations made by Artegiani et al. (1997a), was retained. For the computation of the solar radiation the formula suggested by Rosati and Miyakoda (1988) with the empirical formula of Reed (1977) for the attenuation by clouds was used. The longwave back radiation was calculated from the Bigniami et al., (1995) formula, while the computation of the sensible and the latent heat flux followed the neutral Budyko scheme (Budyko, 1963). For the estimation of the fresh water budget, the evaporation rate (E) was calculated by the latent heat flux (E=Qe/Lv, where Lv=2.65  106 is the latent heat for vaporization) and the precipitation (P) data were taken from Jaeger’s (1976) monthly data set with 5
 2.5
resolution. The river runoff data were taken from Raicich (1994). The river

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runoff (R) is the most important component of the fresh water budget in the Adriatic basin, especially in its northern part, where Po river is the major fresh water source accounting for 30% of the total annual mean basin runoff. For the inclusion of river runoff in the model calculations we follow the parameterization described by Kourafalou et al. (1996a). Parameterization of all major rivers along the Italian and Albanian coast, as well as diffusive sources along the Croatian coast, were included in the model. The mean annual river runoff is 4920 m3/s of fresh water, which corresponds to 1.1 m/yr. River runoff data from Raicich (1994) are believed to be overestimated especially in the NA (N. Pinardi, personal communication). Therefore a flux correction term was added in the salinity flux Ws ; so that surface salinity values were maintained closer to the climatological values. Thus the boundary condition for the salinity at the surface becomes:

KH

 
@S DsH  S S ; ¼ WS þ @z g

where KH is the .vertical diffusivity, Ws the salinity flux (E2P)S; DsH the surface layer depth, g the relaxation time scale, S  the climatological surface salinity and S the salinity of the surface s-layer as it is estimated by the model. The salinity flux correction term is proportional to the depth of the surface s-layer, which in turn is a fraction of the total depth. Therefore the relaxation is strong for shallow areas like the NA (relaxation time scale: B2 days), and almost negligible at the deep SA (relaxation time scale B40–50 days). The estimated mean-annual fresh water budget (E–P–R) gives an annual fresh water gain of 0.6 m.

Table 1 The bulk formulae used for the calculation of the heat fluxes Heat budget components

Bulk formulae

Solar radiation Qs Longwave back radiation Qb Latent heat flux Qe and sensible heat flux Qh

Rosati and Miyakoda (1988) Bigniami et al. (1995) Neutral Budyko (1963)

3. Numerical experiments, results and discussion After a 3 years spin up of the model with the climatological forcing, it exhibited an almost repeated seasonal cycle. We will further down discuss the results of the simulated 4th year.

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3.1. Water masses and circulation patterns 3.1.1. Seasonal variability of the thermohaline circulation The model reproduces the main features of the seasonal circulation of the basin (Figs. 4–7), in agreement with observations and previous numerical studies. The general surface circulation is cyclonic with smaller scale cyclonic features embedded in the global cyclonic pattern. Maximum mean-seasonal surface speeds vary from 0.30 to 0.35 m/s in conformity with the findings of Poulain (2001). In particular, one of the permanent features of the basin is the SA cyclonic gyre, which is controlled by the topography of SA Pit and exhibits a seasonally varying strength. We find cyclonic recirculation around SAP stronger during autumn and winter as a result of the increased buoyancy losses of this period (Fig. 8), while wider

cyclonic circulation in SA remains intense during both winter and spring. The latter intensification is attributed to the intensification of water exchanges through the Otranto Strait following the DWF period. According to Artegiani et al. (1997b), the SA gyre is mostly prominent during autumn and almost absent during winter, while Poulain (2001), in drifter data between 1990 and 1999, finds cyclonic recirculation around SAP intensified in winter and spring. Current measurements for the years 1994–1995 (Kovacevic et al., 1999) show that SA gyre appears stronger in autumn and winter and weaker in spring. The permanent cyclonic circulation in the SA is further confirmed from the numerical simulations of Zavatarelli et al. (2002). In contrast, Gacic et al. (1997), from the analysis of SST fields in the Adriatic Sea from AVHRR data (1984–1992), indicate that the SA gyre is a non-permanent feature of the upper thermocline, but a recurrent one, that occurs only during

Fig. 4. Mean seasonal sy and flow field at 10 m in autumn (Oct.–Nov.–Dec.).

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Fig. 5. Mean seasonal sy and flow field at 10 m in winter (Jan.–Feb.–Mar.).

autumn. Model results are in good agreement with observations (Artegiani et al., 1997b; Poulain, 2001) on the seasonal variability of the bathymetry controlled MA cyclonic gyre, which is found strengthened during spring, summer and autumn. NA circulation is also cyclonic throughout the year. Nevertheless a number of small scale features, like the NA cyclonic gyre present in observations (Artegiani et al., 1997b; Poulain, 2001), are not evident in the model results. As a dilution basin, in terms of its fresh water budget, the Adriatic basin exports low-salinity, riverine water through the Otranto Strait in the surface layer. The offshore baroclinic pressure gradient, induced by the highly buoyant fresh water input in the NA, drives the southeastward current (Western Adriatic Coastal Current— WACC) that flows along the Italian coast of the basin to the Strait of Otranto (Gacic et al., 1996). This current is well developed during spring

summer and autumn in agreement with observations (Artegiani et al., 1997b; Poulain, 2001) and numerical simulations (Zavatarelli et al., 2002). It is weakened during winter by the compensation between temperature and salinity in the density field (Zavatarelli et al., 1999), and it becomes baroclinically unstable during summer. The heat losses at the surface of the Adriatic are compensated by the intrusion of warmer water from the Ionian Sea in the surface and intermediate layers of the Otranto Strait that dominates its eastern part. The inflow of warmer and saltier Ionian Water at the surface layer of the Strait forms the Eastern Adriatic Coastal Current (EACC). In the model results this current, flows up to the northern sub-basin where it re-circulates to meet the WACC (Poulain, 2001). It is mostly prominent during autumn, winter and spring while it is almost absent during summer, because of the upwelling induced by the Etesian winds blowing

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Fig. 6. Mean seasonal sy and flow field at 10 m in spring (Apr.–May–Jun.).

over the region in conformity with observational data (Artegiani et al., 1997b). The thermohaline circulation of the basin is also very much influenced by the river runoff particularly during autumn and spring. A strong salinity front at the surface between the river discharge areas (mainly the Northern and the Southeastern Adriatic) and the rest of the basin throughout the year is evident in the model results.

3.1.2. Heat and fresh water budgets of the basin In our simulations the estimated total heat budget of the climatological year is –24 W/m2, which is close to the estimates of mean annual heat losses made with climatological data of 19–22 W/ m2 by Artegiani et al. (1997a) and 19 W/m2 by Zavatarelli et al. (2002). In Table 2 the monthly averages of the different components of the heat budget are shown.

The fresh water budget (evaporation minus precipitation and river runoff (E  ðP þ RÞ) monthly averages are also shown in Table 2 and give an annual gain of 0.6 m/yr (E  P ¼ 0:5 m/yr, R ¼ 1:1 m/yr). Zavatarelli et al. (2002) estimate a fresh water gain equal to 0.85 m/yr with ECMWF reanalysis data. Artegiani et al. (1997a) estimate 1.14 m/yr for (E  ðP þ RÞ), but they suggest a possible coastal bias and overestimation of their precipitation budget (1.01 m) up to 50% due to the lack of open sea measurements. Our estimates compare very well with the SOC (Southampton Oceanography Center) climatology based on COADS data (Josey et al., 1999) that gives an E2P annual budget of 0.51 m/yr for the Adriatic basin. In addition the estimates of fresh water budget from Raicich (1996) (from 0.65 to 1.10 (70:18) m/yr) and Zore-Armanda (1969a) (from 0.48 to 0.56 m/yr) are also close to our estimates. The fresh water gain is maximun in

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Fig. 7. Mean seasonal sy and flow field at 10 m in summer (Jul.–Aug.–Sep.).

April (1.13 m/yr) and minimum in September (0.13 m/yr). 3.1.3. Seasonal variability in Otranto Strait The mean annual inflow/outflow rate at the Strait in the model results varies from 2.4 Sv in early spring to B1 Sv during autumn with a mean value of 1.6 Sv (Fig. 9a), which is much larger than the estimation of Vetrano et al. (1999) and Gacic et al. (2002), based on current measurements in the Otranto strait during the year 1995, of a mean annual inflow/outflow rate of B1.05 Sv and with a different seasonal cycle that gives a maximum volume transport in autumn and minimum in early spring. However, observational data (Manca et al., 2002) and numerical simulations of Samuel et al. (1999) find that water flux in Otranto strait varies from year to year as a function of the intensity of the Adriatic DWF processes. This is also confirmed from the interannual variability experi-

ments of this study (see Section 3.3). Thus longer time series of current measurements in the Otranto strait that will reveal the importance of the interannual variability of the water fluxes are necessary. The annual heat loss of the basin at surface is offset by the import of the same amount of heat on an annual basis at both the surface and intermediate layer through the Otranto Strait. This heat transport is maximum during spring—in agreement with the maximum volume transport—and minimum during summer (Fig. 9b). An important seasonal cycle is evident in the vertical structure of the water mass exchanges through the Otranto Strait. This seasonal variability depends on the variability of the thermohaline fluxes between the two adjacent basins (Adriatic and Ionian) and the seasonal variability of the wind field and atmospheric pressure gradient (Gacic et al., 2002).

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Fig. 8. Mean buoyancy flux (in m2/sec3) for the Northern Adriatic (), Middle Adriatic (dashed line), Southern Adriatic (solid line).

The outflow of the Adriatic deep water (ADW) occupies the western part of the bottom of the Strait (500–800 m) during all seasons (Fig. 10) as also reported by Vetrano et al. (1999), and it is maximum in March and minimum in autumn (Fig. 9c). The major presence of this water mass in winter and spring, reported by Manca and Giorgetti (1998) and Manca et al. (2002), is in conformity with model results. This water is cold and saline (TB13.2
C, SB38.63) with sy over 29.15. At intermediate layers (150–400 m) Modified Levantine Intermediate Water (MLIW) enters the Adriatic on the eastern part of the Strait with T ¼ 14
C and S>38.7. The inflow of MLIW is stronger during summer and autumn (Fig. 9d) in agreement with observations (Manca and Giorgetti, 1998). At the surface layer the Ionian waters enter on the eastern side and the low salinity Adriatic waters outflow on the western side. The portion of

the strait occupied by the inflow or the outflow at surface varies seasonally. The inflow dominates during winter and spring, while the outflow dominates during summer, in agreement with observations (Orlic et al., 1992). 3.2. Deep-water formation (DWF) Both types of DWF, on the continental shelf and in the open ocean (open ocean convection), that take place in the Adriatic basin, are quite well simulated in our reference numerical experiment. These two processes will be described and discussed separately for all three sub-basins. 3.2.1. Northern and middle Adriatic In the model results, the winter meteorological conditions in the NA sub-basin, with the very dry and cold Bora wind acting in the area during this period, induce the appearance of very low

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Table 2 The surface heat fluxes interactively computed by the model using air temperature, relative humidity, wind, and cloud data from ECMWF ‘79–’94 climatology and the sea surface temperature from the numerical calculations. Qs: incident solar radiation in W/m2; Qb: backward radiation flux in W/m2; Qh: sensible heat flux in W/m2; Qe: latent heat flux in W/m2; Qtot: tolal heat flux in W/m2 and E– P–R: Fresh water budget in m/yr, where E:evaporation rate, P:precipitation rate, R: river runoff Month

Qs

Qb

Qh

Qe

Qtot

E–P–R

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec Mean annual

73.60 115.46 186.14 259.64 325.83 360.91 350.02 302.36 224.16 143.22 87.07 61.66 207.51

114.83 109.15 101.08 99.16 101.54 102.13 96.75 95.67 102.44 110.28 117.58 117.62 105.68

47.22 33.69 9.67 4.64 10.29 13.67 10.31 11.40 18.63 30.69 46.20 53.17 24.13

113.72 97.92 67.13 60.36 72.31 94.88 104.97 105.93 110.76 128.47 137.00 131.16 102.05

202.17 125.31 8.26 95.48 141.69 150.23 138.00 89.36 7.66 126.21 213.71 240.29 24.36

0.82 0.86 1.08 1.13 0.87 0.37 0.19 0.22 0.13 0.50 0.87 1.00 0.6

temperatures. Although the salinity of the area, which is very much influenced by the river runoff, is also low, because of the shallow bathymetry of this sub-basin, the penetration of the atmospheric information to the bottom is not inhibited. The maintenance of high buoyancy losses (0.2– 1.3  107 m2/s3) for an extended time period (from September to February) in the area (Fig. 8) triggers the appearance of very high densities. The DWF period starts in mid January—at least 1 month earlier than the respective period in the SA (Fig. 11). The whole water column is vigorously mixed and ventilated. The density maximum for the whole sub-basin occurs in the middle of February, and the dense water is formed at the central and the northern part of the NA. The sy of the core of NADW exceeds 29.4, with TB9
C and S=38.11. These characteristics are in close agreement with climatological estimates of Artegiani et al. (1997a) with temperature of NADW found slightly lower in the model results. The Middle Adriatic is known as a storage basin for the deep water coming from the NA while local DWF is also reported (Zore-Armanda, 1963). In our results, the continental Bora wind, acting over the region of MA, induces intense buoyancy losses (0.5–1.5  107 m2/s3 from October to March) (Fig. 8) and the MLIW is present at the

intermediate layers (50–150 m). The combination of these two factors favor DWF on the eastern continental shelf during February (with sy > 29:4) and open ocean convection over the Pits that reaches almost the bottom (maximum sy B29:3; TB11:2
C, SB38:3) in the beginning of March. A couple of weeks later the bottom layers of the MA Pits are filled with denser waters originating from the NA (Fig. 11). MADW characteristics are within error limits of climatological estimates of Artegiani et al. (1997a). 3.2.2. Southern Adriatic (SA) In the SA, as in the MA, two distinct dense water masses are identified inside the sub-basin. The first one is locally produced by open-ocean convection over the SAP, in agreement with observations (Ovchinnikov, 1985; Artegiani et al., 1997b), and reaches a depth of approximately 700 m (SADW). The second denser water mass, below 700 m to the bottom of the SAP, is advected there from the northern parts of the basin. This double mechanism provides the overall Adriatic Deep Water (ADW) that exits the basin through the Otranto strait, and it is described in the following two sections. (A) Locally produced deep water: Following the MEDOC Group (1970), we divide the DWF

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Fig. 9. Mean monthly values of: (a) volume transport (inflow/outflow) (Sv), (b) total heat transport (Watt), (c) Adriatic Deep Water (ADW) outflow (Sv) and (d) Modified Levantine Intermediate Water (MLIW) inflow (Sv) at the Otranto Strait (ADW: (sy > 29:15) and LIW: (S > 38:7)).

processes in the SA, into three phases: (a) the preconditioning phase, (b) the violent mixing phase and (c) the sinking and spreading phase. (a) The preconditioning phase: The preconditioning mechanisms in the SA, as revealed in the model results, include: (1) the existence throughout the year of a weak stratification in the center of the cyclonic circulation (SA gyre) trapped over the local topography of the SA Pit, which results in the doming and outcropping of the isopycnals near the surface. This cyclonic gyre has a diameter of B150 km. The erosion of the thermocline in the center of the SA gyre is evident even in periods of strong stratification. (2) the existence of the characteristically saline MLIW at intermediate depths. The intrusion of the MLIW in the basin between 150 and 400 m is stronger during summer

and autumn (Fig. 9d). The major portion of this water mass (more than 90%) re-circulates in the SA and participates in the forthcoming DWF process. When the mixed layer reaches the maximum salinity layer its density is rapidly increased. (b) Violent mixing phase: This phase includes two different processes, the convective process and the baroclinic adjustment process, which interact and finally define the characteristics of the deepwater formed. (b1) Convective process: During winter, the intense atmospheric forcing over the preconditioned area of SA induces great buoyancy losses (Fig. 8) in the area. The horizontal extent of such forcing is larger than the DWF area, and its maximum occurs at the northeastern part of the SA, where the northeastern continental Bora wind

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Fig. 10. sy transects and normal velocilty at Otranto Strait (200-800 m) in (a) autumn, (b) winter, (c) spring, (d) summer.

is more intensive. As a result of the interaction of this forcing with the topography and the existing circulation, a second smaller cyclone inside the larger one, trapped by topography, appears in the SA Pit. This inner gyre has a diameter of B80 km and it is the area of the open ocean convection over the SA Pit. Thus, in agreement with Lascaratos (1993) and Alverson and Owens (1996), the topography selects the location and sets the horizontal scale for the chimney forma-

tion. In particular, this distinct inner cyclone is evident from the beginning of February, when the atmospheric forcing induces the outcropping of the isopycnals at depths greater than 400 m ðsy > 29:15) and permits the ventilation of the waters in the narrow bottom layer of the SA Pit. This inner cyclone is asymmetrically embedded in the larger one, bounded by the 1000 m isobath. The two cyclones share their eastern flow and are distinct at their western part. The bifurcation

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Fig. 11. Deep-water formation rates (in Sv) in the climatological year of (a) the Northern Adriatic Deep Water (NADW), (b) the Middle Adriatic Deep Water (MADW) and (c) the Southern Adriatic Deep Water (SADW). (—): sy > 29:1; (--): sy :> 29:15; (- -): sy > 29:2; (-  -): sy > 29:25; (..): sy > 29:3:

occurs at the northernmost edge of the two cyclones, where the topography drives the inner cyclone to the west. The topographically trapped fluid inside the inner cyclone re-circulates within this limited area of the Pit, and thus it is exposed

to the atmospheric forcing for longer periods. Hence, this limited area becomes the location where the buoyancy forcing is the most efficient at generating strong vertical mixing and becomes the formation site of the SADW. The vertical

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Fig. 12. Meridional velocity cross-section at Southern Adriatic Pit on 15th March.

structure of both gyres in middle March (Fig. 12) shows that the external gyre is stronger than the inner one as a consequence of the larger horizontal density gradients at its edges and the inner gyre has a stronger barotropic component. The maximum mixed-layer depth is B700 m, and the maximum sy of this water mass is 29.18, achieved in the middle of March (Fig. 16). The temperature and salinity characteristics of this water mass are 13.3
C and 38.67 respectively and the estimated formation rate of the deep-water formed inside the Pit with sy between 29.15 and 29.2 is 0.31 Sv. These results are in agreement with climatological estimations (Artegiani et al., 1997a,b). Experimental studies (Gacic et al., 1998; Manca et al., 2002) show that the intensity of vertical convection in SA varies greatly on an interranual basis. There are winters when the convective mixing is practically absent (1993–1994,1996–1997) and others when convective mixing reaches the bottom (1986–1987, 1991–1992). In 1999 maximum mixed layer depth inside SAP is found to be 700 m (Manca et al.,

2002). A comparison of the sy profiles in middle February 1999 (not shown) with the model results (Fig. 17) shows similar structures. Dense water with sy > 29:2 occupies depths greater than 800 m; less dense water, with sy o29:12; is on the eastern side of the SAP, and more dense water coming from the north, with sy > 29:2; enters the SAP on the western continental slope. However, the depth and density of the mixed layer at this period as well as the width of the convected area are larger in the model results. (b2) Baroclinic adjustment process: The external gyre becomes unstable and meanders develop at the periphery of the gyre almost simultaneously with the onset of the convection process, in compliance with the theoretical predictions of Legg et al. (1998) who suggest that the necessary condition for the baroclinic instability to occur is provided by the convection, and it is the reduction of the Rossby deformation radius of the basin. The physical mechanism of the amplification of these meanders is a release of potential energy—initially

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Fig. 13. Snapshot of sy and flow field at 30 m in the Southern Adriatic (a) on 1st March and (b) on 15th April (reference run).

increased by the thermohaline forcing—through a baroclinic instability process (Madec et al., 1991; Jones and Marshal, 1993).

The meanders have a wavelength of L ¼ 2pR where R is the Rossby radius of deformation for the 1st baroclinic mode. In the model results

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L ¼ B30 km in the first days of February. The wavelength of the meanders does not increase with time as suggested by Madec et al. (1991). Mesoscale anticyclonic and cyclonic eddies are more evident at the northwestern part of the external gyre compared to its eastern part and induce inhomogeneities in the density field (Fig. 13a). This northwestern part of the gyre is the site of a density front between denser water (with lower temperature and salinity) coming from the north and lighter water (with higher temperature and salinity) originating from the external cyclonic gyre. A possible explanation, for the minor presence of eddies on the eastern side of the gyre, is that buoyancy losses there are larger than those on the western side. According to Killworth (1976) and Madec et al. (1991), buoyancy loss input can swamp the development of eddies and promote the convective process. Moreover Madec et al. (1996) and Legg and Marshall (1998) find that the development of the baroclinic instability of the convected region can be suppressed from an ambient circulation, comprising a barotropic gyre. Thus the further development of the baroclinic instabilities, especially at the inner gyre, may be inhibited by the strong barotropic component of both gyres (Fig. 12), which is also confirmed from observations (Kocacevic et al., 1999). Several previous studies on water mass formation (Legg and Marshall, 1993; Herman and Owens, 1993; Send and Marshall, 1995; Visbeck et al., 1996; Lascaratos and Nittis, 1998) have shown that the role of these eddies is to transfer buoyancy laterally from the periphery to the center of the gyre and therefore limit the depth of convection. These lateral fluxes of buoyancy, if large enough, can theoretically completely offset the surface forcing and produce slant-wise convection (Legg and Marshall, 1993). Laboratory experiments of Visbeck et al. (1996) have shown that the time required to reach this quasiequilibrium state (teddy) is equal to 12 (r2/B)1/3 , where r is the radius and B the buoyancy loss of the convection region. In all these studies though, a sharp front between the convected region and the exterior flow exists and the less buoyant water occupies the

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interior of the gyre. In our simulations, contrary to the previously described case, the periphery of the gyre, where the instabilities are present, is dominated by less buoyant water than the interior, coming from the north. Cyclonic eddies bring less buoyant water towards the gyre and anticyclonic eddies transfer buoyant water outside it. Thus these eddies rather diminish this density front at the periphery of the gyre than limit the convection depth inside it. The lateral advection of water induced by this small number of eddies at the periphery of the SA gyre towards its center seems to promote or at least not affect the convective event over the SAP. Assuming that the deepening of the mixed layer in the convection region is one dimensional and there is no entrainment of stratified fluid from the base of the mixed layer, the time required for convection to break through the stratification (tbreak ) is calculated by the equation: dh/dt=B0/ N2h (Turner, 1973) where h is the mixed layer depth, B0 the buoyancy loss and N the Brunt Vaisala frequency. If tbreak is smaller than the previously defined teddy ; then eddies do not have the time to reach finite amplitude and control the convection process through limiting the deepening of the mixed layer. The equation is solved numerically using a vertical resolution of 1 m and a time step of 1 h for both the exterior and the interior gyre. The initial stratification of the beginning of winter and the mean winter buoyancy flux of the 4th year (steady state) are used for both gyres. For the outer gyre r ¼ 150 km, the mean mixed layer depth 200 m, B0 ¼ 8:5  108 m2/s3 tbreak ¼ 15 days and teddy ¼ 89 days, while for the inner gyre r ¼ 80 km, the mean mixed layer depth 500 m, B0 ¼ 9  108 m2/s3, tbreak ¼ 49 days and teddy ¼ 57 days. Therefore tbreak is smaller than teddy and the eddies do not have the time to take control on the convection process and limit drastically the mixed layer depth especially for the outer gyre. In order to further quantify the role of mesoscale eddies in our experiments we conduct a one-dimensional(1-d) experiment—where eddies are of course absent—using a 1-d version of POM. The initial conditions for this experiment are provided by the T and S profile of a point inside

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Fig. 14. (a) Initial sy profile of 1-d experiment, (b) mixed layer depth at the end of February (solid line: 1-d experiment, dashed line: 3-d experiment).

the SAP on 31st December (4th year of the 3-d 1 experiment with horizontal resolution of 50 th of a degree). Starting from there, the model is forced for 3 months with identical atmospheric forcing as in the 3-d experiment. The initial profile and the depths of the mixed layer for the 1-d and 3-d experiments for the end of February are presented in Fig. 14. The comparison of the mixed layer depth shows that this is mainly determined by the vertical convection induced by the surface atmospheric forcing, while the lateral buoyancy transfer by eddies offsets a relatively small part of the surface buoyancy loss and affects the finally attained sy : The influence of the mesoscale eddies on the DWF rates is also examined by comparing the 1 results of a lower resolution (10 th of a degree) 1 simulation and a finer resolution (50 th of a degree) 1 with those of the reference one (20th of a degree). The higher resolution the more destabilized the density field is due to the development of

baroclinic instability (Fig. 15). Nevertheless the 1 th and difference between the DWF rates of the 10 1 th grid runs is negligible, while there is a 10% 20 1 decrease in DWF rate of the 50 th grid run. Thus the contribution of mesoscale eddies to the lateral buoyancy flux, during the convection process, is presumably small. This result applies to the Adriatic and cannot be generalized to the whole Mediterranean. It is interesting to note that Lascaratos and Nittis (1998) using POM model in the Levantine basin found teddy quite smaller than tbreak suggesting that the eddies in the LIW formation site play a significant role in the laterally buoyancy transfer in the convected area. In conclusion, it is likely that the overall effect of the baroclinic adjustment process is less important and the convective process is predominant in SAP. However, as observational data show, the interannual variability of buoyancy forcing plays a crucial role in determining the

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Fig. 15. Snapshotss of sy and flow field at 30 m in the Southern Adriatic on 1st March for (a) the 10 km grid run and (b) the 2 km grid run.

scales of convection mixing and eddy activity in the area. Moreover numerical experiments forced with higher temporal and spatial resolution of

atmospheric forcing are required to reveal a detailed overall picture of the competition between the two processes.

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(c) Sinking and spreading phase: By the middle of March, the heat losses at the surface cease and the vertical mixing stops. Then, geostrophic adjustment takes place and the restratification at the surface starts with buoyancy input through the air–sea fluxes. This process is further facilitated and accelerated by the presence of the baroclinic instabilities that bring buoyant water towards the convected region (Fig. 13b). The efficiency of this process depends on the stratification of the ambient flow. Its timescale is estimated from the idealized experiments of Jones and Marshall (1997) as trestrat ¼ 56 r (h Db)1/2, where r is the radius of the convected area, h is the depth of the stratified layer adjacent to the convection region and Db ¼ gDr=r0 the density jump across it. At the eastern part of the gyre, where initially the stratification of the ambient flow is stronger, the meanders develop faster into mesoscale eddies that transfer buoyancy towards the gyre, shrink the collapsed chimney of dense water and eliminate the density anomaly. Assuming a stratified layer of 400 m depth and Db ¼ 1:7  103 m/s2 (corresponding to a 0.18 sy difference) at the eastern side of the SA gyre, the estimated restratification timescale is trestrat B62 days, which is in close agreement with model results. Fig. 17 shows that by the middle of May (2 months after the end of the convection period) isopycnal 29.18 has reached 400 m. In the same figure it is evident that restratification is faster at the eastern side of the gyre, where the stratification of the ambient flow is stronger due to the intensification of EACC during this period that brings buoyant Ionian surface water (warmer and less saline) inside the Adriatic basin. At the same time, at depth (not shown), mesoscale cyclonic eddies trap dense water inside them and carry it away from the convected region. The temporal evolution of the dense water volume in the Adriatic (divided by the time period of 1 year in order to transform it to transport units) in each sub-basin during a whole climatological year is shown in Fig. 11. In contrast to the NA and MA it is obvious that there is always a background water of high density (sy > 29:15) occupying the SA Pit throughout the year. The strong seasonality in the estimated DWF rate is

linked to the seasonal outflow regime of this dense water through Otranto (Figs. 9c and 11). The maximum volume of water with sy > 29:15 present in the SAP is observed in March. The duration of the deep-water production period in SA is about 1.5 months (from the beginning of February to the middle of March), while the deep-water spreading period away from SAP towards Otranto strait, is much longer (almost 10 months: from the middle of March to the end of December). (B) Deep-water advected from the north: Dense waters formed in the NA and the MA propagate south as a vein of dense-water flowing along the western shelf isobaths in accordance with the conservation of potential vorticity for frictionless motions (Bigniami et al., 1990a,b), Zoccolotti and Salusti, 1987). This vein of dense water is characterized from low temperature and salinity. When it reaches the vicinity of the city of Bari (Fig. 1), a topographic anomaly, as also reported by Bigniami et al. (1990a,b) and Manca et al. (2002), destabilizes the flow and the dense-water sinks and fills the bottom of the SA Pit (Fig. 16). This dense-water mass contributes to the dense water outflowing through the Otranto Strait (ADW). The vein fills the bottom of the SAP from March until August with dense-water formed in nearby areas coming first and dense-water from the NA coming last. As shown in Fig. 17 the isopycnal 29.2 inside the SAP, which characterizes the water advected from northern regions, is uplifted from B800 m in mid February to B650 m in mid-May, and it is found at B700 m in mid-August (not shown). The annual mean rate of the deep-water formed in the northern parts of the basin is estimated, at its source, to be 0.1 Sv (B0.033 Sv for NA and B0.067 Sv for MA and western continental slope of SA) with high sy ranging from 29.3 to 29.9. This dense water, while advected from its source towards southern regions, loses its original characteristics and it is detected at the bottom of the SAP with sy between 29.2 and 29.3 with a slightly lower annual mean rate of B0.07 Sv. Our estimates are close to the estimates of Artegiani et al. (1997b) who find a DWF rate ranging between 0.007 Sv and 0.07 Sv for NA, and between 0.04 and 0.3 Sv for SA (sy > 29:1). However, their estimated

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Fig. 16. sy transects at the Southern Adriatic Pit (SAP) on 15th March. Note that the mixed layer exceeds 700 m at the center of the ! near Bari SAP ðsy ¼ 29:18) and the vein of deep-water on the western shelf propagates southwards and fills the bottom of the SAP (transect c).

rate for MA (between 0.002 and 0.007 Sv for MA (sy > 29:2)) is much smaller than our estimation. They also speculate that a contribution to the deep-water pool of the SA could come from MA and NA deep-waters. In conclusion, inside the SAP, the estimated DWF rate with open ocean convection and sy

ranging between 29.15 and 29.2 is 0.31 Sv (0–700 m), while the amount advected from northern region dense water with sy between 29.2 and 29.3 is 0.07 Sv (700–1200 m). Thus we conclude that the dense water present in the SAP has sy ranging from 29.15 to 29.3 and 18% of this dense water originates from the northern parts of the basin.

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Fig. 17. sy transect at the Southern Adriatic Pit (transect c of Fig. 13) in (a) mid-February, (b) mid-March, (c) mid-April and (d) midMay. Note that the depth of sy ¼ 29:2 (shown with the arrow) is uplifted from B800 m in mid February to B650 m in mid May.

3.2.3. Outflow of the deep water through Otranto Strait The deep-water formed in the Adriatic Sea (ADW) exits through Otranto Strait into the Ionian Sea. In Fig. 10a transect at Otranto Strait with the seasonal cycle of ADW outflow is presented. This outflow has the form of a bottom density-driven current, occupies the deep part of

the western continental slope of the Strait between 500 and 800 m in agreement with observations (Vertrano et al., 1999; Gacic et al., 2002), and is stronger from the middle of March until the end of June. In this period the layer of the deep water at the Strait is thick and extends up to 300 m at the western part of the Strait. The temperature and salinity characteristics of the deep water are

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13.2
C and 38.64 respectively, with sy > 29:15: The outflow of the deep water is still strongly evident during summer but its volume has been decreased. In the presence of strong stratification during autumn the remaining deep water of the basin continues to exit occupying a thin bottom layer. This outflow is now weaker and the water is lighter due to its mixing with the surrounding waters during its long residence time in the basin. In early winter even though the mixed layer has already started to increase the deep water, being formed at this period inside the basin, has not reached the strait yet. The estimated deep-water outflow rate with sy > 29:15 is B0.28 Sv in compliance with previous estimations (Lascaratos, 1993; Roether and Schlitzer, 1991; Gacic et al., 1996). Therefore, from the 0.38 Sv of deep water with sy > 29:15 produced and advected in the SAP, 0.1 Sv recirculates within the basin, mixes with lighter water and loses its characteristics.

3.3. Sensitivity of deep-water formation rates and characteristics to interannual variability of forcing factors In order to examine the impact of an interannually variable heat and river runoff forcing on the DWF rates of the Adriatic basin, a number of sensitivity experiments were performed, using the 1 coarse horizontal resolution model grid (10 th). These experiments started from climatology, then ‘‘extreme’’ forcing was applied followed by ‘‘normal’’ climatological forcing (see details below). We studied the effect of this abnormal year on the

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DWF rates, as well as its effect, if any, on the DWF rates of the following ‘‘normal’’ years. The experiments, four in number (Table 3), were designed in the following manner: from the 3rd year (steady state) of the climatological runs, where we used the interactive scheme for the estimation of the heat fluxes, we diagnosed the mean monthly heat fluxes and salinity fluxes, which were then imposed to the model for 1 year (no salinity flux correction term was used) (Fig. 18, years: 0.75–1.75). We then conducted a 1-year run with ‘‘non-climatological’’ forcing (Fig. 18, years: 1.75–2.75). In this year we mimicked a ‘‘cold’’ (exp.1) or a ‘‘warm’’ (exp.2) year, by subtracting or adding a mean annual heat flux of 30 W/m2, to the imposed climatological heat forcing. This ‘‘non-climatological’’ heat forcing in experiments 1 and 2 was imposed only during the cooling period (namely from October to March) so as to affect only one convective cycle. The mean value of 30 W/m2 per year falls within the estimated variability of the atmospheric forcing in the area (Maggiore et al., (1997)). Alternatively we mimicked a year with half (exp.3) or double (exp.4) Po river-runoff (compared to the climatological Po river runoff). After this ‘‘non-climatological year’’, we ran the model for two more years with the climatological forcing (Fig. 18, years: 2.75–4.75). The results of exp.1 (Fig. 18) show, as expected, that the DWF rate, in all areas are increased during the ‘‘cold’’ year (months 19–30), and higher densities are also produced. It is interesting to note that during the next climatological year we still observe higher than ‘‘normal’’ DWF rates. This

Table 3 The experiments 1–4 for the sensitivity of the DWF rates on interannual variability of the heat and river runoff forcing Years: 0.75–1.75 (1st Oct–30th Sept) Exp.1 Exp.2 Exp.3 Exp.4

Climatological runoff Climatological runoff Climatological runoff Climatological runoff

heat fluxes and river

Years: 1.75–2.75 (1st Oct–30th Sept) 2

heat fluxes and river

Decreased heat fluxes by 30 W/m (mean annual) Increased heat fluxes by 30 W/ m2(mean annual) Half Po river runoff

heat fluxes and river

Double Po river runoff

heat fluxes and river

Years: 2.75–4.75 (1st Oct–30th Sept) Climatological runoff Climatological runoff Climatological runoff Climatological runoff

heat fluxes and river heat fluxes and river heat fluxes and river heat fluxes and river

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Fig. 18. DWF rates (in Sv) in the sensitivity experiments 1 and 2 of the Northern Adriatic Deep Water (NADW), the Middle Adriatic Deep Water (MADW) and the Southern Adriatic Deep Water (SADW). The ‘‘extreme’’ heat forcing is imposed during the years 1.75– 2.25 (6 months). 1, (--): sy :> 29:15; (—): sy > 29:2; (-  -): sy > 29:25; (..): sy > 29:3:

implies the existence of a ‘‘memory’’ mechanism of the ocean to the atmospheric conditions of the preceding year. In other words the DWF rates in the basin depend not only on the heat forcing of the present year but on the time history of the heat

forcing as well. During the second climatological year, DWF formation rates are almost restored for the shallow MA and NA while denser water, in comparison to that of climatological years, is still produced in SA.

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The results of exp.2 (Fig. 18) support the previously defined ‘‘memory’’ of the basin, which is still valid in the climatological year(s) following the ‘‘warm’’ year (decreased heat losses). In all three basins, the DWF event is almost absent during the ‘‘warm’’ year. In the SA, the volume of high-density water (sy > 29:15), present in the SAP from the previous climatological year, decreases during this period, as it continues to exit the basin through the Otranto Strait, without being replaced by newly formed waters. No deep water is formed in NA during the two ‘‘normal’’ years following the ‘‘warm’’ one and DWF is very weak in MA and SA as well. The relative importance of the period of the year that the ‘‘abnormal’’ forcing is imposed is also examined. Two additional experiments (not shown) were conducted with the same annual mean heat loss for the ‘‘abnormal’’ year, as in exp.1 The heat loss is now imposed during the whole year (exp1.1) or only during the warm period (April–September) (exp1.2). The results of these experiments reveal that the ‘‘memory’’ is still evident. In exp1.1 the increased heat losses affect two convective periods and thus the duration of the memory is longer. In exp.1.2 the memory effect is less evident due to the less effective downward transfer of the atmospheric information during the stratification period. The ‘‘memory’’ is different for each sub-basin as a consequence of the differences in their bathymetry and their proximity to the open boundary (Otranto Strait). Moreover the intensity of water fluxes in Otranto follows the intensity of the DWF processes inside the basin in agreement with observations (Manca et al., 2002). Its duration can not be easily estimated and certainly depends on the intensity of the buoyancy forcing as well as the dimensions of the sub-basins. Among the mechanisms that can possibly explain the ‘‘memory’’ of the system in these experiments we can include: (a) the thermal preconditioning of the basin produced by the presence of colder/warmer waters inside the basin following the cold/warm ‘‘extreme’’ atmospheric year (locally produced or advected though the open boundary); (b) the salinity preconditioning produced by the increased/decreased intrusion of salty surface and

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intermediate waters through the Otranto Strait as a response to increased/decreased dense water outflow. A detailed evaluation of the relative importance of the two above mechanisms is beyond the scope of this study. Moreover these experiments are schematic and take into account the interannual variability of only one of the components of the buoyancy forcing (heat forcing), and thus quantitative conclusions would be unrealistic. However, these experiments clearly show that a ‘‘cold’’ cooling period in 1 year is a positive preconditioning for the DWF processes of the next year(s), while a ‘‘warm’’ cooling period is a negative preconditioning for the DWF processes of the next year(s). The results in experiments 3 and 4 show that the decreased/increased runoff of the Po river mainly affects the DWF processes in the NA and not the open ocean convection over the SA Pit, where the influence of the riverine waters is less important. In particular exp.3 shows an increase of the volume of deep-water formed in NA, while the increase of Po runoff in exp.4 decreases the area of DWF in NA, which is now almost all occupied by less saline riverine waters. In both experiments water exchanges between the sub-basins are increased but through a different mechanism in each experiment. In exp.3 the mechanism responsible is the increased deep-water volume, while in exp.4 it is the increased buoyancy input at the surface. Gacic et al. (1996) suggest that an increased runoff intensifies the overall cyclonic circulation of the basin and thus strengthens of the intrusion of the surface Ionian Water and of MLIW at the intermediate layer that potentially influences positively the DWF processes, particularly in the Southern sub-basin. This is not confirmed in our results since the intensification of the cyclonic circulation induced by the increased Po river runoff does not seem to influence the DWF processes in the SA.

4. Discussion and conclusions The 3-d primitive equation POM model with a horizontal resolution of 1/20th of a degree was implemented in the Adriatic basin in order to study the general thermohaline circulation of the

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basin with particular focus on the processes related to DWF occurring in its sub-basins during winter. A climatology of 16 years (1979–1994) ECMWF Reanalysis data (1
 1
resolution) was used for the atmospheric forcing with particular attention to the construction of the wind field and wind stress climatology, so that the intensity and the main directions of the wind were well represented. The heat fluxes were not imposed but interactively estimated by the model, using the sea surface temperature values produced by the model itself. The major river discharges were also parameterized inside the model giving an annual mean river runoff of 4920 m3/s. The results of the simulated 4th year forced with the climatology show that the major features and their seasonality are correctly simulated by the model in accordance with observations and previous numerical studies (Orlic et al., 1992; Artegiani et al., 1997b; Poulain, 2001; Zavatarelli et al., 2002). The estimated annual mean heat budget is – 24 W/m2 and the annual-mean fresh water budget is 0.6 m/yr (fresh water gain). The basin scale circulation is mainly cyclonic, while permanent features such as the SA cyclonic gyre exhibit an important seasonal variability in their strength and dimensions. The main water masses of NADW, MADW and SADW, with T and S characteristics similar to those observed, are correctly formed by the model during a climatological year. The Adriatic Deep Water (ADW) that exits the basin through the Otranto Strait is the deep-water concentrating inside the SAP with two different mechanisms. The upper 700 m of the SAP are occupied by the SADW, which is locally produced by open-ocean deep convection inside the Pit. The bottom layer of the SAP is filled with deep-water advected from the northern areas of the basin, where it is originally formed mainly on the continental shelf. The NADW is formed during February on the continental shelf of the NA sub-basin with sy > 29:4: Deep water with high sy is also formed on the continental shelf of MA and local open-ocean deep convection over the MA pit occurs during winter. Those two dense water masses are temporarily stored in the MA Pits and then flow southwards as a vein of dense water between 50

and 150 m along the western shelf isobaths, in accordance with observations (Zoccolotti and Salusti, 1987; Bigniami et al., 1990a,b). The dense water of the vein sinks, mixes with the surrounding waters and fills the bottom of the SAP from March to August, as soon as it finds a topographic anomaly near Bari. The annual rate of the deepwater formed in northern regions and advected in the SAP is estimated inside the Pit to be 0.07 Sv with sy between 29.2 and 29.3. The SADW is formed over the SAP with openocean convection, inside a second smaller topography driven cyclone, embedded in the larger one which covers the whole SA- between the middle of February and the middle of March. The mixed layer inside the chimney exceeds 700 m and the s–y of the deep-water formed is 29.18. The annual SADW formation rate in the SAP with sy between 29.15 and 29.2 is estimated to be 0.31 Sv. Baroclinic instabilities appear at the SA gyre from the beginning of the convective process in agreement with theoretical predictions of Legg et al. (1998). The development of baroclinic eddies at the periphery of the gyre, usually expected to be important (Legg and Marshall, 1993), is found limited presumably due to: (a) the intensity of the buoyancy forcing, (b) the barotropicity of the ambient cyclonic circulation (Legg and Marshall, 1998), especially at the eastern part of the gyre, and (c) the existence of dense water (denser than the interior of the gyre) at the northwestern part of the periphery of the gyre. The suggestion that mesoscale variability does not affect importantly the convection process in the SA is further confirmed by: (i) the estimated larger teddy (time required for the mesoscale eddies to develop and control the convection process) in comparison to tbreak (time required for 1-d non-penetrative deepening of mixed layer to break through the stratification), (ii) the comparison between a 1-d and a 3-d model integration showing that the mixed layer depth is mainly determined by the vertical convection induced by the surface atmospheric forcing and (iii) the small differences in the comparison of the DWF rates of our reference eddy-resolving grid run, with the results of lower 1 resolution (10 th of a degree) run and a higher 1 resolution run (50 th of degree). However, the

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competition of the two processes (convection and baroclinic adjustment) is highly sensitive to the interannual variability of the buoyancy forcing as well as to smaller scale processes that can not be resolved with the existing temporal and spatial resolution of atmospheric data used to force the model. Thus more observational data and detailed numerical experiments are important in order to reach definite conclusions. The development of mesoscale eddies though, during the sinking and spreading phase, is the mechanism responsible for the ensuing spreading of the deep water away from the convected region, and the restratification timescale predicted in theoretical studies is confirmed from the model results. From the estimated deep-water volume present inside the SAP, with sy between 29.15 and 29.3, it can be concluded that 18% of this water comes from the northern parts of the basin. The volume of the dense water at the Otranto Strait varies seasonally, being maximum from the middle of March until the end of June and minimum in February. The estimated annual deep-water outflow rate through the Otranto strait, with sy > 29:15; is 0.28 Sv, in agreement with previous estimates (Roether and Schlitzer, 1991; Lascaratos, 1993), while 0.1 Sv of deep-water re-circulates within the basin and loses its characteristics through mixing with lighter waters. Four different sensitivity experiments were performed to reveal the sensitivity of the DWF rates of the basin to interannual variability of the heat and the river runoff forcing. The results of experiments 1 and 2 (see also Table 3) show that the sensitivity of the DWF to the heat forcing variability is important, in agreement with field observations (Gacic et al., 2002), and that DWF rates are very much influenced, namely they increase or decrease following the respective increase or decrease of the heat losses of the basin. These rates remain high or low respectively, even when the heat forcing returns back to the climatology, suggesting that the Adriatic basin has a strong ‘‘memory’’ in its density structure not only of the present year heat forcing but also of the heat forcing of the preceding year(s). Therefore 2 years with the same heat forcing do not necessarily have the same impact on the DWF rates. The

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results of experiments 3 and 4 (interannual variability of the Po runoff) show that Po river runoff mainly affects the DWF processes in the NA and has no important impact on the DWF processes in the SA. The climatological runs performed in this study using perpetual year forcing revealed the basic mechanisms of the DWF processes in the area and the results are well compared with available observational data. They also allowed us to establish the concept of the memory of the Adriatic basin. Obviously the climatological forcing does not allow the episodic nature of the DWF process to be simulated and hence we are now working on interannual runs with 6-h forcing. Future studies of DWF in the Adriatic should also consider the drastic changes in the deep thermohaline circulation of the Eastern Mediterranean (Eastern Mediterranean Transient) during the late 1980s–early 1990s, the effect of this change on the DWF processes in the Adriatic basin, and the interplay between the SADW and the Aegean deep water that now occupies the bottom layer of the Eastern Mediterranean.

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