Chapter 5.7 Mediterranean water and global circulation

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5.7 Mediterranean Water and Global Circulation Julio Candela

5.7.1 Marginal seas Marginal seas bordering the world's oceans make important contributions to the global thermohaline ocean circulation through their exchange of water with the major oceanic basins. Their principal effect is related to processes of water mass formation that occur within these bordering seas where heat and/or net freshwater losses are intensified due to their confined nature. It is well known that some of the intermediate and all of the deep water masses of the world's oceans are produced as the result of intense air-sea exchange in marginal seas (Warren, 1981a). The water masses formed in the marginal sea typically enter the open ocean as a dense outflow through a restricted channel or strait. One of the best known outflows is that from the Mediterranean Sea, which enters the North Atlantic through the Strait of Gibraltar as the lower layer of a two-way exchange flow (Bryden et al., 1994; Bower et al.9 1997). After rounding Cape St Vincent, at the southwestern corner of the Iberian Peninsula, the outflow water begins to spread into the eastern North Atlantic to form a warm and salty tongue that extends westward from Portugal across the eastern Atlantic (Wiist and Defant, 1936; Worthington, 1976; Lozier et al., 1995). This tongue is one of the most prominent features of the North Atlantic hydrography at intermediate depths, and its high-salinity water has been implicated in the preconditioning of the North Atlantic Deep Water formation (Reid, 1978). This chapter deals with the Mediterranean Sea and its influence on the world's ocean thermohaline circulation and therefore on climate. It

OCEAN CIRCULATION AND CLIMATE ISBN 0-12-641351-7

describes water formation processes within the Mediterranean Sea, outlining the characteristics of well-known water formation sites and the evidence of recent shifts in water formation locations within the sea. Then it focuses on the characteristics of the exchange flows between the Mediterranean and the North Atlantic Ocean through the Strait of Gibraltar based on historical as well as recent (1994-96) flow measurements that reveal significant seasonal and interannual variability of the exchange flows, and in particular of the Mediterranean outflow into the North Atlantic. This outflow is then the subject of a detailed description since it determines the way the Mediterranean Water mixes with the surrounding waters in the North Atlantic, creating a warm and saline tongue of water that can be identified throughout the whole North Atlantic at a depth of about 1100m (Fig. 5.7.1). It is fundamental for numerical models investigating the influence of the Mediterranean Water on the thermohaline circulation of the World Ocean to reproduce or parameterize the mixing processes in the outflow accurately, since they determine the properties and the depth of penetration of the Mediterranean Water in the North Atlantic. The chapter ends with a discussion of observational hydrographic evidence, mainly put forth in Reid's (1994) study, which supports the hypothesis that the Mediterranean contribution to the Nordic and Labrador Seas allows the formation of deep waters in the North Atlantic, or at least that without this contribution the waters formed would not be dense enough to penetrate to the depths they presently do.

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60°N

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Fig. 5.7.1 Salinity (a) and temperature (b) contours at 1100m depth in the North Atlantic from Levitus (1982) climatology.The contour interval for salinity is 0.2%o and for temperature is 1°C.

The Mediterranean Sea, the Mare Nostrum of the Romans, is a semi-enclosed sea with a broad history intimately related to the development of western civilization, but only recently have oceanographers recognized that the waters of the Mediterranean provide them with a model of the world ocean itself (Lacombe, 1990). Geographically it

has a zonal extent of about 4000 km and a mean meridional width of 1000 km, with a mean depth of 1500 m and a maximum depth of up to 5000 m in the Ionian Sea. It is divided by the Strait of Sicily into Western and Eastern basins (Fig. 5.7.2). Mainly an evaporative basin, it acts to transform relatively fresh North Atlantic surface water (salinity of

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Mediterranean Water and Global Circulation

45°N [~

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Fig. 5.7.2 Map of the Mediterranean Sea indicating names of places mentioned in the text.The shaded areas in the Gulf of Lions, Adriatic, Aegean and Levantine correspond to known regions of water mass formation.The 200,1000, 2000 and 4000 m depth contours are also shown.

36.1%©) into a dense Mediterranean Water, which has a salinity of 38.4%o, a temperature of 13°C and a density of 1029.7kgm~ 3 , as it outflows at depth through the Strait of Gibraltar. The Mediterranean Sea is often regarded as a 'laboratory basin' for oceanographic circulation studies, especially of the thermohaline circulation, which is particularly active despite the relatively small size of the sea. It is one of the few places in the world where deep convection and water mass formation take place. In the present climate, deep convection occurs only in the Atlantic Ocean; the Labrador, Greenland and Mediterranean Seas; and occasionally also in the Weddell Sea. Convection in these regions feeds the thermohaline circulation, the global meridional-overturning circulation of the ocean responsible for roughly half of the poleward heat transport demanded of the atmosphere-ocean system (Marshall and Schott, 1999; see also Bryden and Imawaki, Chapter 6.1). It is conjectured that the resulting Mediterranean outflow plays an important though indirect role in the North Atlantic circulation (Reid, 1979) and, consequently, in the thermohaline conveyor belt at global scales and on time scales of global climate change (Wu and Haines, 1996). According to Reid (1979) the Mediterranean outflow helps maintain the high salinity of the Norwegian Sea. Without this source of high-salinity water the Norwegian-Greenland

Sea might not provide the denser waters that fill the Arctic Basin and thus contribute a major component of the North Atlantic Deep Water. It is the relatively saline North Atlantic Deep Water, transported by the deep western boundary current, that penetrates into the low-salinity waters of the Weddell Sea, where it is cooled further and enriched with brine to provide the Antarctic Bottom Water - the densest water found in the oceans of the world. To the extent that these suggested linkages are correct, the exchange between the Atlantic Ocean and the Mediterranean Sea is of significant importance. 5.7.2 Formation of Mediterranean W a t e r There are several places within the Mediterranean where preconditioning (i.e. a cyclonic circulation with convex curvature of isopycnals that bring dense, and usually weakly stratified, waters close to the surface) and air-sea fluxes combine to induce convective processes (for details on the convection mechanisms, see Marshall and Schott, 1999). These are the Gulf of Lions in the Western Mediterranean, the region south of Rhodes, the Levantine Basin, the southern part of the Adriatic Sea, and in recent years the south region of the Aegean Sea (Roether et aL, 1996) in the Eastern Mediterranean (Fig. 5.7.2). Of these sites, the Gulf of Lions region

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is where convective processes reaching depths of more than 2000 m have been extensively documented since the Mediterranean Ocean Convective (MEDOC) experiment (MEDOC Group, 1970). The water mass distribution in the Western Mediterranean comprises three layers. In the upper layer, and originating mostly from the inflow through the Strait of Gibraltar, is the Modified Atlantic Water (MAW). Between 150 to 500 m depth, a warm and salty layer is found, referred to as Levantine Intermediate Water (LIW), which is formed by shallow convection in the Eastern Mediterranean Basin and then slowly spreads into the Western Mediterranean through the Strait of Sicily. Beneath the LIW layer, the basin is filled with near-homogeneous Western Mediterranean Deep Water (WMDW). It is this WMDW that is formed by deep convection processes in the Gulf of Lions and which contributes importantly to the characteristic of the Mediterranean outflow (Stommel et al.9 1973; Kinder and Parrilla, 1987). Based on a newly composed hydrographic climatology, Krahmann (1997) estimates a WMDW production rate in the northwestern Mediterranean of 1.8 + 0.6 x 1013m3 yr _1 , corresponding to 0.6 +0.2 Sv. The average yearly WMDW production is made up of 1.3 x 10 13 m 3 of LIW and 0.5 x 1013m3 of MAW. Thus, the formation rate of LIW and slight variations of its characteristics have implications on the WMDW formation rate and its variability. Actually, it is becoming clear from recent observations that deep water formation is not a process that recurs every year with certainty and regularity. The intensity of convection shows great variability from one year to the next and from one decade to another. For example, in the Gulf of Lions, 1969, the year of the first MEDOC experiment (MEDOC Group, 1970), was a year of strong convection, but convection in 1971 was not as strong (Gascard, 1978). Vigorous deep convection to 2200 m returned in 1987 causing a very homogeneous water body of potential temperature 12.79°C and salinity 38.45%o (Leaman and Schott, 1991). Convection reached only to 1700 m in 1991 and did not mix the water column as thoroughly (Schott et aL, 1996). Fluctuations in the composition and possibly also the volume of the Mediterranean outflow are the result of the variability in WMDW formation, coupled to the LIW formation variability (Nittis and Lascaratos, 1998). Recent findings that in the last decade an

influx of Aegean Sea water has replaced 20% of the deep waters of the Eastern Mediterranean, which were previously only formed in the Adriatic Sea (Roether et al., 1996), also affect such fluctuations. Apart from these observed interannual variations in deep and intermediate water formation, there is a well-documented increase in the salinity of the deep waters of the Western Mediterranean over the past 40 years (Lacombe et ai, 1985; Leaman and Schott, 1991). Observations also suggest that in the past 10 years there has been a jump in the salinity of the newly formed deep waters in the Eastern basin (Roether et al., 1996). It has been argued that this salinity increase has resulted from the diversion of the Nile and Russian rivers for irrigation so that the effective net evaporation over the Mediterranean basin has increased. Application of hydraulic control models then project that the overall Mediterranean salinity will increase by about 0.13%o over the next 100 years or so (Rohling and Bryden, 1992). However, it remains unclear what the implications of such an increase in salinity would be on the overall circulation of the Mediterranean and North Atlantic. 5.7.3 Outflow of Mediterranean W a t e r at the Strait of Gibraltar

The Strait of Gibraltar is the Mediterranean's only communication with the World Ocean; it is about 60 km long, 15 km wide at its narrowest section (Tarifa narrows) and only 280 m deep at its main sill (Fig. 5.7.3). There are three main components to the flow (Lacombe and Richez, 1982; Candela, 1991): a tidal, mainly barotropic flow, with magnitudes of up to 2.5ms" 1 (Candela et #/., 1990); a barotropic subinertial component driven by atmospheric pressure fluctuations within the Mediterranean and with magnitudes close to 0 . 4 m s - 1 (Candela et al., 1989); and a baroclinic subinertial component driven by the internal pressure gradient due to the density difference between the Mediterranean and the Atlantic Waters, with magnitudes of about 0.5 ms" 1 , and likely to be hydraulically controlled (Armi and Farmer, 1988). Therefore, the Strait is dynamically very energetic with tidal, subinertial and long-term currents all being of significant amplitude. This situation makes studying the exchange flows particularly difficult, requiring long and careful measurements

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36.2°N

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Longitude Fig. 5.7.3 Map of the Strait of Gibraltar showing the location of the sill mooring indicated in the text (large dot). The 50,100,200,300,400 and 500 m depth contours are also shown. Depths larger than 400 m are shaded.The small dots distributed along Gibraltar's main sill indicate the positions where current profiles were obtained from a ship-mounted ADCP during a tidal cycle.

not only of currents but also of water characteristics and in particular of the fluctuations of the interface that separates the inflowing Atlantic from the outflowing Mediterranean Waters (Bryden £**/., 1994). The Strait of Gibraltar has been the subject of several field measuring programmes in recent years (Bryden and Kinder, 1991). The longest continuous record, 2 years (October 1994-October 1996) of continuous measurements of the current profile and water properties at a mid-sill location on Gibraltar's main sill (Fig. 5.7.3), comes from a 2.5-year measurement programme that concluded in October 1996. These observations were obtained with an upward-looking, bottom-mounted, broadband 150 kHz Acoustic Doppler Current Profiler (ADCP), capable of measuring the entire 280 m water column at this mid-sill location, with a vertical resolution of 10 m. In addition, during two cruises on board the RV Poseidon in April 1996 and in October 1997, consecutive crossings of the strait were performed over the sill section measuring the current profile using a ship-mounted ADCP through a complete semidiurnal tidal cycle (Figure 5.7.3 shows the location of the current profiles measured by the ship over the sill). The April 1996 cruise coincided with a period of neap tides (small amplitude), while those of October 1997 were performed during spring tides (large amplitude), providing an idea of the across-strait

current structure during both tidal extremes. From these sections it is clear that the currents at the sill present large cross-strait variability; however, the mid-sill moored ADCP measurements capture the main time variability of the currents and both sets of observations are used here to estimate a time series of the exchange through the Strait. In calculating the exchange through the Strait, it is important to obtain estimates of the quality, as well as the quantity, of the water being exchanged. In order to distinguish Atlantic from Mediterranean waters it is essential to have simultaneous measurements of the density structure of the water column along with those of the currents. In addition, it is mandatory to take into account the contribution to the mean exchange from the high correlation between the barotropic (tidal and subinertial) currents and the depth of the interface separating Atlantic and Mediterranean Water types at the sill (Bryden et aL, 1994). For this reason, simultaneously with the bottom-mounted ADCP measurements, an additional mooring was installed that contained several (3 to 5) instruments in the water column, depending on the deployment period, which made it possible to construct time series of the depth of the interface between the Atlantic and Mediterranean Water cores. Based on previous work (Bryden et aL, 1994), as well as these observations, it was decided to use the 37 salinity as the characteristic value delimiting the boundary between

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the two layers. Estimates of Atlantic and Mediterranean Water exchanges were calculated using hourly time series of current velocities at 10 m depth intervals from the surface to the bottom, hourly time series of the depth of the interface, and a realistic cross-section bottom relief together with the cross-strait current structure based on the aforementioned ship surveys done during spring and neap tidal cycles. These, after being low-pass filtered to retain periods longer than 3 months, are shown in Figure 5.7.4. An important result from these calculations is that both the Atlantic and Mediterranean Water transports show a small, but appreciable, seasonal cycle. Of the two, the outward Mediterranean lower-layer flow is a more reliable estimate, showing an annual transport range of 0.28 Sv with I

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minimum outflow around early summer (July 1995) and a maximum in late winter (February 1995 and 1996). The seasonal cycle on the inflow is not as well represented in these observations, although maximum inflow tends to occur during the summer of 1995 with a second maximum at the beginning of 1996. Minimum inflows occur around late winter (February 1995 and 1996) coinciding with the maximal outflows. Bormans et al. (1986) suggested a seasonal cycle in the inflow of about 6%, with maximum transport in the spring. They attributed the increase to changes in interface depth and argued that winter water mass formation processes raise the interface level within the Mediterranean, while draining of the Levantine Intermediate Water reservoir occurs during the rest of the year, effectively lowering the interface. The

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Fig. 5.7.4 Low-passed (periods larger than three months) transport estimates in the Strait of Gibraltar.The transport calculations are done from hourly time series of currents measured at the mid-point of Gibraltar's main sill with a vertical resolution of 10 m, time series of the depth of the 37%o isoline separating Atlantic and Mediterranean Waters and information of the cross-strait current structure based on shipboard current observations across the sill during two complete semidiurnal tidal cycles at springs and neaps.The inflow (upper), outflow (middle) and total (bottom) transports are indicated in Sverdrups (1 Sv = 10 6 m 3 s _1 ).The mean, maximum, minimum and range of each plot are also indicated.The time axis of the lower plot is common to all three panels.

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Mediterranean Water and Global Circulation

observations shown in Fig. 5.7.4 do not seem to support this argument. Originally the calculated barotropic or total transport showed a mean value of 0.086 Sv, which is about twice of the expected value based on estimates of net evaporation over the Mediterranean Basin (Garrett et aL, 1993b). This pointed to uncertainties in the estimates of the upper-layer inflow, where the cross-section is wider and transport estimates based on only midstrait current measurements were considered unreliable. Therefore currents in the upper 100 m were reduced by 9% in order to have a mean total transport value of 0.04 Sv over the 2 years of observations, in accordance with historical net evaporation estimates. Apart from this mean correction the seasonal cycle present in the total flow has the correct phase to be the principal contribution to the observed seasonal sea-level rise within the Sea, although to explain fully the observed sealevel change one has to take into account also the effects due to evaporation and seasonal changes in heat content. Evaporation and heat loss from the Mediterranean also have seasonal cycles, with maximum loss of both heat and moisture during winter; heat is actually gained from the atmosphere by the Mediterranean during April-August (Bunker, 1976b). The effect of seasonal air-sea forcing over the Mediterranean on the instantaneous transport through the strait, however, may be quite small, as the residence time involved in the transformation of Atlantic Water to Mediterranean Water is on the order of decades (Lacombe et aL, 1981). It is more likely that seasonal fluctuations of transport result from dynamic effects more local to the strait, or from mechanisms of draining and filling of the LIW and/or WMDW reservoirs, as suggested by Bormans et aL (1986). At periodicities of days to months there is a clear increase of variability of the exchange in both layers during the late fall and winter periods (Astraldi et aL, 1999). At interannual time scales, the formation of Mediterranean Water masses may occur under different conditions in different years, changing the characteristic temperatures and salinities of the outflow (Lacombe et aL, 1985). As a result, different volumes of outflow in different years might transport the same flux properties. This variability in the characteristics of the outflow is generally thought to be a small effect, but investigators have noted a wide range in the maximum of salinity

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observed west of the sill: Schott (1928) found a maximum of 37.25%o over several years of observations, Boyum (1963) a maximum of 38.2%o, and Kinder and Parrilla (1987) a maximum of 38.44%o. Based on hydrographic sections in the Gulf of Cadiz, Ochoa and Bray (1991) estimated an equivalent transport of 'pure' Mediterranean Water of 0.7Sv, assuming salinities of 38.4%o and 35.6%o for Mediterranean and Atlantic Waters. They also estimated a freshwater flux equivalent to 0.53 m y r - 1 of evaporation excess over the Mediterranean and a heat flux higher in the autumn of 1986 than in the spring (6.0 W m - 2 versus 2.2 W m - 2 ) , due to warmer surface temperatures in the autumn. Salt fluxes from our observations, computed by the same method as Bryden et aL (1994), show a mean value of 1.3 Sv%© with a large standard deviation of 0.4Sv%o in the 2 years, which implies a mean net evaporation over the Mediterranean Sea of 0.45myr - 1 , when mass and salt conservation equations for the whole Mediterranean Basin are taken into account. This net evaporation is similar to previous estimates (Bryden et aL, 1994), but shows appreciable interannual differences of 0.15myr - 1 . Temperature fluxes, computed as in Macdonald et aL (1994), give a mean inward flux of 3 Sv °C, with a considerable interannual difference of 1 Sv°C. This mean temperature flux through Gibraltar implies a mean heat loss through the Mediterranean Sea surface of about 3 W m - 2 , considering that the estimated 0.04 Sv of net inward flow leaves the surface of the Mediterranean at 25°C. This heat loss is a factor of two smaller than the 6 W m - 2 obtained by Macdonald et aL (1994) based on direct current and temperature measurements done in the strait between October 1985 and October 1986. However, in contrast, climatological estimates of the basin mean heat flux based on sea surface measurements have typically shown a heat gain by the Sea of 2 0 - 3 0 W m - 2 (Garrett et aL, 1993b; Josey et aL, 1997). Further research is required to resolve this discrepancy. Although both observed salt and heat flux interannual differences seem to be small with respect to their effect on the North Atlantic water mass structure, it is still not possible to conclude that they are negligible. From a climatic point of view it is important to know the mean values of the inflow and outflow as accurately as possible as well as any interannual variabilities. Recent numerical model studies

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(Gerdes et aL, 1999) suggest that a correct representation of the Mediterranean salt tongue in the North Atlantic requires correct inclusion of the outflow mass volume rather than the frequently used procedure of only specifying a salinity source at the location of the Strait of Gibraltar. The transport estimates shown here, which are close to 1 Sv each way, are significantly higher, about 30%, than those calculated from earlier data (Bryden et aL, 1994). This increase principally comes from the mean currents at the mid-sill loca-

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tion being 30% larger in the 1994-96 period from those measured in the October 1985-October 1986 period (Fig. 5.7.5). This significant change in magnitude between two measuring periods separated by nearly a decade is surprising and points to the importance of monitoring the exchange flows at Gibraltar as an indication of long-term climatic changes occurring in the Mediterranean Sea. Another important aspect of these new measurements is that they show relatively large interannual differences of the outflow, increasing from 0.98 Sv

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Fig. 5.7.5 Mean (a) and rms values (b) of the along-strait current measured at the sill in the Strait of Gibraltar. The continuous lines correspond to observations obtained with a bottom-mounted ADCP at the location indicated in Figure 5.7.3.The ADCP profiled the whole water column, with a 10 m depth resolution, at four 6-month-long deployment intervals between October 1994 and October 1996.The specific time interval for each deployment is indicated in the plots. Concurrent with the sill ADCP there were currents measured with Aanderaa current meters on a nearby mooring.These are indicated by large dots at the specific depth of each instrument. Each dot corresponds to a 6-month-long measuring period within the October 1994 to October 1996 observation interval. Also indicated are currents measured with Aanderaa current meters at a nearby sill location (squares and diamonds), but during the Gibraltar Experiment (GIBEX) between October 1985 and October 1986. Details on the specific location of the M2 and F3 moorings, as well as the length of the measuring periods for each instrument during GIBEX, can be found in Candela eta/. (1990).

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in February 1995 to 1.07 Sv in March 1996. This increase of close to 0.1 Sv from 1995 to 1996 is significant and could very well have implications on the effects of the outflow in the water structure of the North Atlantic. Close monitoring of the Gibraltar Exchange will certainly render valuable information for climate studies of the Mediterranean Sea and the Global Circulation. 5.7.4 The effect of Mediterranean W a t e r outflow on the circulation of the North Atlantic and the World Oceans

The Mediterranean Water exits the Strait of Gibraltar as a single, dense plume that flows down the northern continental slope of the Gulf of Cadiz as a gravity-driven boundary current over complicated topography. As the outflow spreads northwestward along the southern Spanish and Portuguese coasts, it slowly loses its high salinity as it mixes with fresh North Atlantic Central Water (Baringer and Price, 1997a; Bower et al., 1997). By the time the flow reaches the vicinity of Cape St Vincent it is neutrally buoyant (Ochoa and Bray, 1991; Zenk and Armi, 1990). From this point, part of the outflow is trapped along the continental slope flowing northward, while another part flows westward into the ocean interior (Reid, 1994; Iorga and Lozier, 1999a). It is known that a large portion of this westward branch is implicated in the formation of sub-mesoscale coherent vortices, Meddies, that contain a core of warm and salty Mediterranean Water (Kase and Zenk, 1996; Bower et al, 1997). Meddies are typically 20-100 km in diameter and 200-1000 m thick, centred about 1000 m depth. Richardson et al. (1989) estimated that 8-12 Meddies form each year based on estimates of the number of coexisting Meddies and their average lifetime (2-3 years), and that they may transport about 25% of the salinity anomaly flux that comes through the Strait of Gibraltar. Arhan et al. (1994) suggested that Meddies may be responsible for more than 50% of the zonal salinity flux at the level of the Mediterranean Water in the North Atlantic, based on simultaneous observations of three Meddies along a hydrographic section at 15°W, and previous work showing that some Meddies drift westward at a speed five times the background flow. Recently, Bower et al. (1997) have estimated a Meddy formation rate of 15-20 Meddies per year

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based on observations of RAFOS (reverse of SOund Fixing And Ranging) floats seeded in the Gibraltar outflow in the Gulf of Cadiz. This Meddy production rate supports the idea that a large portion of the observed zonal salinity anomaly flux, about 50%, is directly related to Meddies. With regard to the observational basis for the pathways of the Mediterranean outflow, we do not yet know how much of the outflow proceeds north along the Iberian continental slope and how much flows or circulates slowly westward from the Strait. We cannot yet estimate with certainty the outflow advected as Meddies versus how much is advected in concentrated boundary currents, just as we are still unclear on the mixing mechanisms that erode the high-salinity core of the outflow (Bryden and Webb, 1998). Quantifying and understanding these details are at the heart of being able to answer with certainty the role played by the Mediterranean outflow on the North Atlantic circulation. The outflow of the Mediterranean is only about 1 Sv. This is a relatively small transport compared with those found in the North Atlantic, but its salinity of 38.4%o and temperature of 13°C are extremely high compared with any other waters in that depth range. The effect of these large contrasts in water mass properties is two-fold: first, they help identify both a northward flow along the eastern boundary to the Greenland-Scotland sill and a westward flow across the Atlantic that turns southward along the western boundary, reaching the Antarctic Circumpolar Current and the Weddell Sea. Second, even at those distant places where this water shallows or outcrops, such as near Iceland and in the Weddell Sea, it retains salinities high enough to form, when cooled enough, the densest waters of the northern North Atlantic and the Weddell Sea (Reid, 1994). Also, the Mediterranean outflow is the source of the mid-depth changes of heat and salt to the waters entering the North Atlantic. As the Mediterranean outflow pours down from the Strait of Gibraltar it joins the subsurface waters flowing northward along the eastern boundary. These northward flows include waters from the South Atlantic, and before they interact with the outflow they are relatively low in temperature and salinity. After the encounter, these flows carry the outflow water northward to about 40°N and then divide, a portion continuing northward along the boundary

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FORMATION AND TRANSPORT OF WATER MASSES

towards Iceland and the rest turning with the westward limbs of the large cyclonic and anticyclonic gyres. The northward part carries warm and saline water that mixes with the NorwegianGreenland Sea Waters in the passages east and west of Iceland. The westward part crosses the Atlantic, and as it approaches the western boundary near 25-35°N it turns both northward, forming the deepest part of the Gulf Stream above about 3000 m, and southward, joining the southward flow along the deep western boundary and the eastward limb of the cyclonic gyre. As a result of these gyral flows, the added heat and salt from the Mediterranean outflow are spread both northward by the cyclonic gyre and southward along the western boundary into the South Atlantic Ocean, where its contribution makes the southward flow warmer and more saline than the incoming circumpolar waters to the east (Reid, 1994). However, recent analysis based on a new North Atlantic climatology (Iorga and Lozier, 1999a,b) does not support the presence of a westward current of Mediterranean Water that crosses the North Atlantic. Instead, Iorga and Lozier's inverse model exhibits a westward current at ~35°N, which derives its waters from the poleward, eastern boundary current off Africa rather than from the Mediterranean outflow. This would imply that the main mechanism for the westward propagation of the Mediterranean outflow is 'diffusion' by meddies rather than a well-defined current. With respect to the northward flow as an eastern boundary current, their work does give evidence for the presence of the northward branch of Mediterranean Water penetrating past 55°N on its way to the Nordic Seas. However, recent work by McCartney and Mauritzen (2001) based on regional water mass distributions and geostrophic shear rejects Reid's deep source hypothesis in favour of a shallow source hypothesis. Rather than a flow of deep Mediterranean Overflow Water along the eastern boundary rising from depth to feed the Nordic Seas, the inflow is supplied directly by transformed North Atlantic Current waters from the same depth range as the inflow. This shallow source hypothesis is also supported by the recent high-resolution ocean circulation modelling work of New et al. (2000a), which proposes shallow sources composed of water masses of western origin carried by branches of the North Atlantic Current and the more saline Eastern

North Atlantic Water transported northwards from the Bay of Biscay region in a 'Shelf Edge Current' around the continental margin. Clearly more work is needed to identify if any of the proposed sources of saline waters to the Nordic Seas is exclusive of the others, or more likely the main contributor during certain periods. In any event what is well recognized is that the warm and saline deep water flowing out of the North Atlantic extends throughout the Antarctic, Indian and Pacific Oceans, identified as a salinity maximum that lies above the bottom over most of the southern hemisphere and projects back into the Atlantic Ocean through the Drake Passage. The salinity decreases along the path of flow but it remains high enough where the saline layer shallows in the Weddell Sea to contribute to the formation of the densest abyssal water of the world's open ocean (Reid, 1994). A possible line of research to investigate the implication of the Mediterranean outflow on the North Atlantic circulation is to use numerical simulations, though they must be validated through a systematic comparison of their output with actual observations. Recent modelling efforts have been quite successful at reproducing the observed structure and have helped us to understand the dynamics governing the behaviour of the Mediterranean salt tongue in the North Atlantic (Spall, 1999; Stephens and Marshall, 1999). However, it seems essential that the models should not only simulate the mean observed features, but they should also reproduce the observed transient behaviour of the system. For climate change studies the models should exhibit interannual and longer-term variations of the same nature and magnitude as are observed. Models useful for predicting climate changes resulting from an increase in salinity of the Mediterranean outflow should capture or parameterize accurately critical small-scale processes, such as the pathways of the descending outflow plume after it exits the Strait of Gibraltar and its mixing. This poses a particular challenge for numerical algorithms because the strong current is of such small scale and involves mixing processes that occur on even smaller scales. Hence, it is of outmost importance to establish some degree of confidence in the models' ability to represent the real outflow characteristics by a thorough validation based on comparing the model simulations with real data. Before model error estimates are

5.7

Mediterranean Water and Global Circulation

available, model simulations and predictions must be taken with care. After proper validation, the models can be used as a means of obtaining predictions of the response of the outflow to external forcings. Such experiments may be used also to study feedback processes and to identify the principal elements involved. However, one has to be very careful not to stretch the available evidence and venture into unsupported speculations on what might happen if, for example, the present trends of salinity increase of the deep waters in the Mediterranean persist; and in what ways this might affect the deep water formation processes in the North Atlantic, and consequently the thermohaline circulation of the world's oceans. All evidence points towards a very important contribution of the Mediterranean outflow to the thermohaline circulation of the World Ocean and therefore on climate. Nonetheless, the evidence remains sketchy and incomplete. Much further research is required to enable us to understand the mean state and the variability of the Mediterranean outflow and to model its characteristics properly in large-scale climate models, before a certain assessment can be made. Both components are important because their proper understanding

would provide models with a good chance of correctly predicting future changes in the outflow and their consequences for the ocean's circulation and climate. Acknowledgements

The reported transport measurements in the Strait of Gibraltar between October 1994 and October 1996 were done in collaboration with Richard Limeburner from the Woods Hole Oceanographic Institution in Woods Hole, Massachusetts, USA, and with Juan Rico Palma from the Instituto Hidrografico de la Marina in Cadiz, Spain. These measurements were supported by the US Office of Naval Research contract N00014-94-1-0347 and by grant OCE-93-13645 from the US National Science Foundation. The Instituto Hidrografico de la Marina in Cadiz, Spain, provided invaluable logistic support during the field work and the use of the hydrographic ships B.H. Tofino and B.H. Malaspina. Uwe Send from the Institut fur Meereskunde Regionale Ozeanographie, Kiel, Germany, is greatly acknowledged for providing shipboard ADCP observations collected by the RV Poseidon in April 1996 and October 1997.

Candela

429

Net Surface Heat Flux, 1991-1993 Plate 5.1.9

90°N

(see p. 329) Global distribution of the 1991-93 average net heat flux, <*in>, after I 6 . 4 W m - 2 has been subtracted everywhere to make the global average zero. The contour interval is 25 W m - 2 and positive (red) values indicate ocean heating.

360°

Net Surface H2O Flux, 1991-1993 Plate 5.1.11

90°N

(see p. 330) Global distribution of the 1991-93 average flux of fresh water, <^ i n >.The contour interval is I 0 - 5 kg/m-2 (300 mm y r - ] and positive (red) values indicate an excess of precipitation over evaporation.

contoured at [...,-1,0,1,-] x 10 5kg/m 90°S

2

270°

360°

<%n>

360°

Plate 5.1.13

(see p. 331) Global distribution of the 1991-93 average surface density flux, < 3 i n > , computed from the

fluxes of Figs 5.1.9 and 5.1.1 I (see Plates 5.1.9 and 5.1.1 I).The contour interval is I mg m~ s~ and positive (red) values indicate areas where this flux acts to increase surface density.

Zonal Mean 3€in Anomaly, Indian Ocean 60°N I i | i i i | i i i | i i i | i i i | i i i | i i i | i i i

Zonal Mean 3€in Anomaly, Pacific Ocean

Zonal Mean Win Anomaly, Atlantic Ocean

40° h Contour interval 5 W/m 2

9

Zonal Mean S£jn Anomaly, Pacific Ocean

Zonal Mean ^ i n Anomaly, Indian Ocean

- •-• °1

Zonal Mean 9^in Anomaly, Atlantic Ocean

60°N | i | i i i | i i i | i i i | i i i | i i i | i i i | i i i

40c Contour interval 100 kg/rrryear

60°S 1983

'85

Plate 5.1.16

'87

'89

"91

'93

'95

'97

1983

'85

'87

'89

'91

'93

'95

'97

1983

"85

'87

'89

'91

'93

'95

'97

(see p. 335) Sixteen-year time series ( 1 9 8 2 - 1 9 9 7 ) of annual heat (top panels) and freshwater (bottom

panels) flux anomalies, zonally averaged over the Indian (left panels), Pacific (centre panels) and Atlantic (right panels) Oceans.The contour intervals are 5 W nrT greater ocean surface heating or freshening.

and 100 kg m~

yr~ , respectively, with positive anomalies (in red) indicating

0 500 1000 1500 -p 2000 £ 2500 4 Q_ Q 3000 + »-_

0 to 3500 4000 4500 5000 5500 6000

0.1

0.2

0.3

0.4

0.5

0.6

0.7 0.8 0.9 -4 2 -1

Diffusivity (10

22

m s )

Plate 5.2.5 (see p. 352) A composite vertical section of diapycnal diffusivity based on velocity micrestructure data spanning the Brazil Basin (after Polzin et a/., 1997).The section runs from the continental slope off Brazil east to the MidAtlantic Ridge crest. A non-linear colour scale denotes the diffusivity. Dissipation profile data were depth-averaged over 250 m and, where multiple stations were available, ensemble-averaged. Underway bathymetric data were used to draw the bottom profile. 60°W

80°S

I >300 • Plate 5.4.2

250 - 300

I 200 - 250

150-200

100-150

50-100

0-50

(see p. 376) Depth of the 95% oxygen saturation, using data sets provided by J. Reid and A. Mantyla (personal

communication, 1998) and many of t h e W O C E Hydrographic programme (WHP) sections. Ranges of proxy of mixed-layer depth are shown by dots with different colours and sizes from yellow (smallest) dots showing 0-50 m, to green (largest) dots showing depth greater than 300 m.This oxygen horizon is a rough proxy for winter mixed-layer depth (Reid, 1982); careful treatment of each geographical region's mixed layers would refine this horizon (Reid, !982).AfterTalley (1999a).

80°S

60°W

Plate 5.4.3

0°

60°E

120°E

180°

120°W

(see p. 376)

(a) Mode water distributions in the world's oceans, after Talley (1999a). Red coloured areas show the subtropical mode waters (STMWs) associated with the subtropical western boundary currents in each ocean (the first type). Pink coloured areas show the eastern type of subtropical mode waters (the second type), including Madeira Mode Water, North Pacific Eastern STMW and South Pacific Eastern STMW. Brown coloured areas show the third type of subtropical and subpolar mode waters, including North Atlantic Subpolar Mode Water, Subantarctic Mode Water and N o r t h Pacific Central Mode Water. Approximate potential densities (Go) are indicated. Black arrows denote the subtropical gyre circulation. See the text for explanation for each type of mode water. (b) Low-salinity intermediate water distributions in the world's oceans, after Talley (1999a). Shown are the North Pacific Intermediate Water (light green), Antarctic Intermediate Water (green), overlap of NPIW and A A I W (medium green), and Labrador Sea Water (blue). The location of formation for each intermediate water is shown with an X. Regions of strong mixing near the ventilation sources that strongly affect the characteristics of the new intermediate waters are shown with cross-hatching.

J

1955

Plate 5.4.4

i

i

i

l

i

i

i

i

]~n

i

F?

ij |

i

i

i

i

i

i

i

i

i

i

i

i

i

1965

i

| | i ^i

I

i

iI

II

n iI lj

iI

iI

II

iI

iI

iI

iI

II

iI

rT

iI

rI

i

i

i

I

i

i

i

I

i

i

i

i

I

i

i

i

i

JJ

i

1975

1985

ly

I

1995

(see p. 385) (a) Potential vorticity (PV) at the Bermuda time series station (Panulirus) at 3 2 ° I 0 ' N ,

64°30'W. Blue is potential vorticity less than 0.7,green is 0.7-I.0, and yellow is 1.0-1.3 x 10"

m~ s~ . (b) Salinity on

isopycnals 26.2-26.8 09. Solid curves are low-passed with a cut-off period of 2 years. Dashed curves are the actual data, (c) North Atlantic Oscillation (NAO) index (atmospheric pressure at the Azores minus atmospheric pressure at Iceland, December-February average), courtesy of D. Cayan. Reproduced from Talley (1996b).

1000

Q_

0

Q

2000

3000

100

300

500

700

900

Distance (km) Plate 5.5.4 (see p. 390) Salinity distribution along the WOCE AR7W CTD line (Fig. 5.5.1) obtained between 19 and 23 June 1993.The water between 500 m and 2200 m in the central part of the section is unusually homogeneous because of deep convective mixing during the severe winter of 1992-93.The CTD station positions are indicated by numbered triangles along the surface.

34.70 IT

34.60 L Dec Months, 1994-95

Apr

June

Plate 5.5.5 (see p. 391) Time series of G\.s from five depths at 56.8°N 52.5°W between June 1994 and June 1995.The sudden increase in G1.5 at 260 m and 5 I 0 m in early and mid-February indicate the arrival at these depths of the deepening convection layer.The slow decrease in 0\.s between June 1994 and the arrival of convection in February is associated with the restratification of the water column following homogenization during the 1993-94 winter.

TU91 3.00 2.00 1.00 0.50 0.25

0.10 0.05 0.00

TU

6.0 4.0 2.0 1.0

Q.

§

0.5

600

0.2 0.1 0.0

30°S Plate 5.8.1

20°

10°

(see p. 433) (a) Tritium section along I35°W (WHP section PI7C) occupied in 1991. (b) Same

as (a), except for He (in TU).

21 22 CO

23

O) JX:

24

0

25 26

27 30°S 100

50

20° 30

10° 20

0° 15

(a)

12.5

10° 10

20° 7.5

5

30°N 2.5

Years

220°W

200°

180°

160°

140°

120°

100°

(b) Plate 5.8.4 (see p. 434) Distribution of the tritium/ 3 He age along (a) 135°W (WHP section PI 7C) and (b) 9.5°N (WHP section P4).

0

3

10

20

30

59

90

100

110 142

150

160

170

180

190

200

0 7.00

mm^^^amm. 500;

2.50 1000;

*%

m

M.II

• 2.00

1500;

1.60

2000;

1.20

»2500-"

0.80

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0.60

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^

3500-

0.40

4000;

0.30

4500;

0.20

5000;

0.10

5500;

0.02

CFC-11 along A17 (pmol/kg)

6000 : 500

1000

1500

2000

2500

3000

3500

4000

4500

5000

5500

6000

6500

7000

7500

8000

0.00 8500

Distance (km) 50°S

45°

40°

35°

30°

25°

20°

15°

10°N

10°

Plate 5.8.5 (see p. 436) Vertical distribution of CFC-I I along WOCE section A17 extending through the western South Atlantic Ocean measured in 1994 (Messias and Memery, 1998).

Plate 5.8.8

(see p. 438) Time series of the

tritium/ 3 He age along the WHP section AR7W across the Labrador Sea. For explanation, see text.

1985

1990

Years

1995

50

60

9480 70

20

30

40

6

500-

1000

1500

2000

2500

"5_ 3000 CD

Q 3500

4000

4500

45.86 45§8

45.90

5000

5500

CFC-11 along 52 W (p mo l/kg)

6000 500

1000

1500

2000

2500

3000

3500

4000

Distance (km) —i—i—i—i—i—i—

10°

15°

-i

20°

1

1

1

25°

r-

i

'

30°

•

i

i

i

'

35°

«

i

>

i

i

40°N

Plate 5.8.7 (see p. 437) Vertical distribution of CFC-I I along WOCE section A20 (52°W) in the North Atlantic Ocean measured in 1997 (Smethie, 1999).

i

•

93 97101

105 113 179 109 117

40°S

171

30°

159 150 142 134 126 163 155 146 138 130 121 11098 86 78

20°

28 37 46 7269 65 6157545148 444137 33 29 25211712 11 10 9 8 7 6

10'

40°

40 38 36 34 32 30 28 26 24 22

20 18 16 14 12 10

8

6

4

2

4 2

50°N

0

Plate 5.8.15 (see p. 443) S3He (%) contoured along WOCE Hydrographic Programme section PI 7 from 1000-m depth to the bottom. Contour interval is 2%, precision of the measurements is about l a = 0.2% in 8 He.

160°E

Plate 5.8.17

160°W

120°W

80°W

(see p. 444)

Near-bottom radiocarbon distribution where bottom depth is at least 3500 m. Both WOCE and GEOSECS data were used in preparing this objective map. The flow of Circumpolar Deep Water northward along the

Near Bottom Values A 1 t% 0

western Pacific island arch system and then clockwise north of the equator is strongly implied.

-240

-220

-200

-180

-160

-140

-120