Formation of an Azores Current Due to Mediterranean Overflow in a Modeling Study of the North Atlantic

Transcription

1 2342 JOURNAL OF PHYSICAL OCEANOGRAPHY Formation of an Azores Current Due to Mediterranean Overflow in a Modeling Study of the North Atlantic YANLI JIA Southampton Oceanography Centre, Southampton, United Kingdom (Manuscript received 24 August 1998, in final form 5 November 1999) ABSTRACT A mechanism for the formation of the Azores Current is proposed. On the basis of observations and model results, it is argued that the primary cause of the Azores Current is the water mass transformation associated with the Mediterranean overflow in the Gulf of Cadiz. Observations show that the transport of the Mediterranean outflow water through the Strait of Gibraltar increases significantly as it descends the continental slope by entraining the overlying North Atlantic Central Water. This entrainment process introduces a sink at the eastern boundary to the ocean upper layer in addition to the inflow into the Mediterranean. Such a sink is capable of inducing strong zonal flows such as the Azores Current. This mechanism is confirmed by numerical experiments with and without the representation of the Mediterranean overflow process. The numerical model is based on the Miami Isopycnic Coordinate Ocean Model. The model does not include the Mediterranean overflow explicitly, but restores the model density fields in the Gulf of Cadiz toward the observations. This restoring condition produces a reasonable representation of the water mass transformation deduced from observations. The formation of the Azores Current in response to the water mass transformation in the Gulf of Cadiz suggests that the Mediterranean overflow is not only a source of warm and saline water at depth, but also has a strong dynamic impact on the ocean upper layer. This study emphasizes the need to improve the representation of the Mediterranean overflow process in general circulation models in order to capture the correct characteristics of the flow fields and water masses in the subtropical eastern North Atlantic. 1. Introduction Southeast of the Grand Banks of Newfoundland the Gulf Stream (GS) separates into two branches. The northern branch turns northeastward and becomes the North Atlantic Current (NAC). The southern branch, which becomes the Azores Current (AC), heads southeastward across the Mid-Atlantic Ridge to the south of the Azores, then flows mainly eastward at a latitude of about 35 N to the Gulf of Cadiz (GoC). Associated with the AC is a front with significant temperature and salinity contrasts. There have been a number of detailed hydrographic surveys of the front at various locations, for example, to the southeast of the Grand Banks (Mann 1967, 1972; Clarke et al. 1980), in the region of the Mid-Atlantic Ridge (Gould 1985; Sy 1988; Stramma and Muller 1989), and southeast of the Azores (Käse and Siedler 1982; Käse et al. 1985; Siedler et al. 1985; Rios et al. 1992). Geostrophic transport fields obtained from historical hydrographic data indicate that the eastward flow extends all the way to Corresponding author address: Dr. Yanli Jia, Southampton Oceanography Centre, Empress Dock, Southampton SO14 3ZH, United Kingdom. Yanli.Jia@soc.soton.ac.uk the African coast with southward branches in the Canary Basin as part of the subtropical gyre recirculation (Stramma 1984; Olbers et al. 1985; Klein and Siedler 1989). The hydrographic database of Lozier et al. (1995) reveals a coherent AC that stretches across the eastern half of the basin, with divergences to the south and convergences from the north such that the downstream transport does not change much. Recent hydrographic surveys also indicate the eastward extension of the AC to the Moroccan continental slopes (Fernández and Pingree 1996; Pingree 1997). Buoys deployed in the AC are found to travel eastward and reach the western side of the GoC, and then move northward or southward along the continental slopes. Based on the above surveys, the AC is observed to be a meandering jet km wide with an eastward velocity of cm s 1. The eastward flow is mostly in the upper few hundred meters but can reach as deep as 2000 m. The current carries a large fraction of the water entering the eastern recirculation region of the Canary Basin. The estimates of the AC transport are in the range of Sv (Sv 10 6 m 3 s 1 ). The surface temperature and salinity changes across the front can be as large as 2 C and 0.3 psu. The front marks the northern boundary of the 18 C Sargasso Sea water in the central North Atlantic American Meteorological Society

2 SEPTEMBER 2000 JIA 2343 Both drifter data (Richardson 1983; Krauss and Käse 1994; Brügge 1995) and satellite altimetry (e.g., Le Traon et al. 1990; Wunsch and Stammer 1995; Stammer 1997) show a band of high eddy kinetic energy (EKE) associated with the AC. Käse and Siedler (1982) observed considerable meandering of the front southeast of the Azores with mesoscale eddies on both sides of the front. Baroclinic instability has been identified as one of the mechanisms for the high energy level (Beckmann et al. 1994a). The isopycnic potential vorticity analysis of climatological hydrographic data (Stammer and Woods 1987) indicates that a necessary condition for baroclinic instability, the reversal at depth of the meridional gradient of potential vorticity, is present at the AC. The primary mechanism for the formation and maintenance of the Azores Front (AF) is unknown. For this very reason, numerical modeling of the AF and its associated variability has been difficult. The front is either absent or very weak in general circulation models. For instance, the high resolution model of the North Atlantic developed under the Community Modelling Effort (CME) does not produce the separation of the GS into the NAC and the AC (Bryan and Holland 1989). The model AC develops only in the eastern basin (Spall 1990). The EKE at the AC latitudes is barely above the background level of variability (Treguier 1992). The meridional density gradient associated with the model AC is too weak and the necessary condition for baroclinic instability is not satisfied (Beckmann et al. 1994a). No significant improvement is found with increased horizontal resolution (Beckmann et al. 1994b). This paper reports the occurrence of an AC in a general circulation model of the North Atlantic. A mechanism for the formation of the AC is first proposed. The characteristics of the AC in the model are then presented and compared with what we know from available observations. Further sensitivity experiments are performed to verify the proposed mechanism. 2. A mechanism for the formation of the Azores Current The presence of the AC, a coherent zonal flow in the inner subtropical gyre that stretches across a large extent of the basin, is a major feature of the circulation of the North Atlantic. It is situated well to the south of the mean zero wind stress curl where Ekman pumping implies southward transport in the ocean, thus its zonal orientation cannot be fully explained by Sverdrup dynamics. There must be other mechanisms involved. One interpretation was given by Käse and Krauss (1996) through a close examination of a time series of the wind stress curl over the North Atlantic (monthly mean and zonally averaged between 65 and 5 W). They showed that in the latitudinal band between 35 and 50 N, the wind stress curl, though significantly negative, is highly variable with annual and interannual fluctuations. They argued that under such a variable condition, nonlinear inertial effects may become the dominant dynamic control in the GS extension region, thus the separation of the GS into the NAC and the AC. They also suggested that a local relative minimum in the magnitude of the wind stress curl exists in the inner subtropical region, which permits a quasi-zonal flow to continue to the eastern basin, hence the eastward extension of the AC. In this study, a complementary mechanism for the formation and maintenance of the AC is proposed. It is suggested that the water mass transformation associated with the Mediterranean overflow in the GoC may induce the AC. The streamfunction field for potential density surface in Lozier et al. (1995) clearly shows the convergence of the streamlines associated with the AC in the GoC, which suggests a possible connection between the two. The Mediterranean overflow is often poorly represented in general circulation models, which may be the cause for a nonexistent AC in many cases. The AC in the eastern basin and the GoC occupy a similar latitudinal extent and this may not be coincidental. The following is an argument to explain how this may operate. At the Strait of Gibraltar dense Mediterranean water spills over the sills into the North Atlantic. The transport of the Mediterranean outflow water at the western end of the Strait of Gibraltar is typically 1 Sv (Lacombe and Richez 1982; Bryden et al. 1994; Baringer and Price 1997). Intense mixing in the GoC increases this transport by a factor of about 3 by entraining the overlying North Atlantic Central Water (NACW) (Ambar and Howe 1979; Ochoa and Bray 1991; Baringer and Price 1997). This entrainment process in the GoC introduces a sink at the eastern boundary to the ocean upper layer in addition to the inflow into the Mediterranean. Such a sink is capable of inducing strong zonal flows such as the AC. There have been earlier laboratory experiments and theoretical studies to suggest that horizontal circulation can be deduced from a given distribution of sources and sinks. Such a source (sink) could be a direct injection (extraction) of flow into (out of) the system, or through vertical flux of mass across density surfaces. In a laboratory experiment, Stommel et al. (1958) showed the induction of basin-scale zonal flows by a point source and sink placed near the eastern boundary in a rotating system. In a theoretical analysis, Pedlosky (1996) showed that, for a localized source or sink of finite extent situated some distance away from the lateral boundaries, zonal flows form west of and within the latitudinal band of the source or sink. The zonal flows must be bidirectional under the constraint of zero pressure gradient east of the source or sink and outside the latitudinal band containing the source or sink. The transport of each of the zonal flows is large in relation to the strength of the source or sink, but the difference in the zonal flows is equal to the source or sink. A similar circulation pattern was obtained by Luyten

3 2344 JOURNAL OF PHYSICAL OCEANOGRAPHY TABLE 1. The surface-referenced potential density anomaly ( 0 )of the 20 layers defining the discretised vertical coordinate of the model. Layer 1 is the model mixed layer that has a variable density distribution. Layer 0 Layer Variable TABLE 2. The diffusion velocities (cm s 1 ) for momentum, tracers and layer thickness for the 1 3 and the 4 3 models. Each diffusion coefficient is the product of the appropriate diffusion velocity and the grid spacing. 1 3 model 4 3 model Momentum Thickness Tracers device to simulate the observed transport pattern associated with the Mediterranean overflow. Sensitivity experiments show that this restoring condition is the driving force for the model AC. When this condition is removed, no AC is formed. and Stommel (1986) in a two moving layer, two gyre ocean circulation model where a downward interfacial (buoyancy) flux was applied within a circle of finite radius in the eastern basin but away from the eastern boundary. Dipole circulation forms on both the upper and lower moving layers west of the circle and meridional flows exist only inside the circle. A different flow pattern is predicted by the eastern boundary ventilation theory of Pedlosky (1983). By allowing a downward transfer of mass at the eastern boundary, an eastward flow into the boundary results in the upper layer and a westward flow out of the boundary in the lower layer. The flow is unidirectional within each layer and the transport matches the downward mass transfer. The entrainment process by the sinking of the Mediterranean overflow in the GoC involves a downward mass transfer in a region attached to the eastern boundary. The source (in the lower layer) and the sink (in the upper layer) are finite but not isolated. We may envisage a situation where both the above forcing mechanisms (Luyten and Stommel 1986; Pedlosky 1983) operate in the system and set the upper and lower limits on the strength of the zonal flows west of the GoC. The flow exchange at the Strait of Gibraltar may introduce an additional forcing to the zonal flows. The theoretical model of Webb (1993), which considers a layer of fluid in the ocean interior including horizontal viscosity but neglecting vertical viscosity, suggests that a point source or sink at the eastern boundary results in an east west jet, which becomes wider toward the west. Within the jet, there is a balance between viscosity, which works to spread the jet, and Rossby waves propagating from the eastern boundary, which works to constrain it in a north south direction. The transport of the jet is determined by the strength of the source or sink. In the present modeling study, the Mediterranean overflow process is not explicitly resolved but is represented by restoring the model density field toward observed values in the GoC. It is shown that this restoring condition is capable of extracting lighter water from the upper ocean and replacing it with heavier water at depth, thus serving as a water mass transformation 3. Model description The model used in this study is based on the Miami Isopycnic Coordinate Ocean Model (MICOM) using surface-referenced potential density ( 0 ) as the vertical coordinate. The details of the model numerics can be found in Bleck et al. (1992). The model domain covers the North Atlantic basin from approximately 20 S to 70 N and from 100 Wto20 E. The horizontal resolution is 1 3 in longitude by 1 3 cos( ) in latitude (where is latitude), thus yielding an isotropic horizontal grid. There are 20 layers in the vertical. The top layer is a mixed layer of Kraus Turner formulation where density and other model variables are allowed to vary with time and in space. The 19 layers below the mixed layer are layers with constant potential density. The values of the layer densities (Table 1) are chosen so that the water masses and the associated dynamics in the North Atlantic can be represented as well as possible with the limited resolution. The bathymetry is taken from the ETOPO5 database from the National Geophysical Data Centre; no smoothing is applied but a minimum ocean depth of 75 m is set in the model. Subgrid-scale processes are parameterized in Laplacian form. There are three lateral mixing parameters for isopycnic diffusion of momentum, layer thickness, and tracers. They are written in the form of diffusion velocities and represent the ratio of the diffusion coefficients to the model horizontal grid spacing. These diffusion velocities are kept constant in the model (Table 2), thus diffusion is proportional to grid spacing. The diapycnic mixing coefficient for tracers is a function of stratification. It is inversely proportional to the Brunt Väisälä frequency (N): a 0 /N, where a 0 is set to 10 7 m 2 s 2. There is no diapycnic mixing of momentum between layers. The wind stress, friction velocity (used in computing turbulent kinetic energy for mixed layer forcing) and the heat flux used to force the model are taken from a three-year monthly climatology derived from the ECMWF analysis ( ). The full surface heat flux consists of the ECMWF climatology plus a restoring term toward an equivalent sea surface temperature with

4 SEPTEMBER 2000 JIA 2345 a variable timescale as described in Barnier et al. (1995). The surface salinity is restored toward the Levitus (1982) climatology with a timescale identical to that used in the heat flux formulation. The model boundaries are closed in the north, south, and at the Strait of Gibraltar. Connection with the ocean exterior to the model domain is parameterized by means of buffer zones in which model variables (depth of interfaces below the base of the mixed layer and salinity of the density layers) are restored toward observed values. The restoring of layer interface depth is designed to account for the processes that determine the observed density structure in the buffer zones. The inclusion of salinity restoring on the density layers does not influence the dynamics (as the density of each layer is fixed), but provides the layers with the appropriate water mass characteristics (temperature and salinity). In the version of MICOM used in this study, salinity is the active thermodynamic variable in the density layers, which is advected and diffused. Layer temperature is derived from layer density and salinity. The northern buffer zone covers the region north of 67 N between 40 and 10 W, and north of the line connecting 67 N, 10 W and 60 N, 17.5 E (17.5 E is the model eastern boundary). Data used for the restoring in this region is based on a hydrographic database from the National Oceanographic Data Center (NODC Informal Report 12, 1991). This database provides an improved description of the water mass characteristics in this region. The combination of a relatively large buffer zone and an improved hydrographic database is designed to achieve a realistic representation of water mass transformation occurring north of the ridges connecting Greenland, Iceland, and Scotland. The Levitus (1982) climatology is used for the restoring in the southern boundary buffer zone and in the GoC. The southern buffer zone is defined south of 11.5 S. The Strait of Gibraltar in the model is at 36 N, 6 W and the associated buffer zone is bounded by 33.5 N and 38.0 N, 11 W and 6 W. Full depth restoring is applied in the northern and southern buffer zones. The restoring timescale increases linearly from 3 days at the model boundaries (70 N and 19 S) to 100 days at the inner edges of the buffer zones. In the GoC, the restoring condition applies down to the upper interface of density layer (model layer 15) and to the salinity of the layers above that interface (at approximately 1500-m depth). The timescale increases linearly from 14 days at the Strait of Gibraltar to 100 days at a distance of 300 km. Levitus (1982) September climatology is used to initialize the model except for the region north of 60 N and east of 40 W where, for consistency, the same database as for the northern boundary restoring is used. Initially the North Atlantic basin is mostly occupied by the top 16 layers; the bottom 4 layers reside mainly north of the ridges connecting Greenland, Iceland, and Scotland. The model is then integrated for a total length of 20 years. The model fields are saved every three days for the last five years of the integration and are used to calculate the time mean fields. A coarse resolution ( 4 3 ) version of the model is set up in a similar way. The mixing parameters are given in Table 2. The sizes of the buffer zones are the same as in the 1 3 model. This model is used for sensitivity experiments presented in section 5. The high resolution ( 1 3 ) version of the model was developed and integrated under the DYNAMO (Dynamics of North Atlantic Models) project funded by the European Union. It was used in an extensive modeling intercomparison exercise involving three models. A detailed account of the model intercomparison can be found in DYNAMO Group (1997). 4. Model results: 1 3 model In this section, the characteristics of the AC in the 1 3 model experiment are presented and compared with observations. Both instantaneous and mean fields are shown. Unless stated otherwise, the mean fields are the annual averages over the last five years of the model integration. a. Near-surface circulation The mean circulation pattern at 110 m is depicted in Fig. 1. The major currents of the GS system are all present: the Florida Current, the GS, the NAC, and the AC. The model shows a significant improvement over previous models in its representation of the AC. However, the GS separates from the coast too far north ( 40 N) compared with observations ( 35 N). Consequently, downstream, the separation of the GS into the NAC and the AC is unrealistic. There are two identifiable origins for the AC, one at approximately 42 N, 47 W southeast of the Grand Banks and the other at Cape Hatteras at about 35 N. The latter is not present in observations although it appears to be the major source for the model AC. The former is in agreement with the analysis by Klein and Siedler (1989) based on historical hydrographic data, although the precise position is slightly too far north. According to Klein and Siedler (1989), there is a direct current connecting the source region (approximately 40 N, 45 W) and the AC region southwest of the Azores in winter, while in summer the flow separates from the source region into two branches, one flowing directly towards the AC region and the other taking a cyclonic loop before joining the AC. This loop is present in the model throughout the year with slight strengthening in summer. The direct route from the origin to the AC is missing from the model. The model AC appears as a wide zonal jet in the mean field between 32 and 35 N, 50 and 25 W. The instantaneous field (Fig. 2) exhibits a much more vigorous flow pattern with a tighter jet (150 km) and large

5 2346 JOURNAL OF PHYSICAL OCEANOGRAPHY FIG. 1. The mean velocity at 110 m of the 1 3 model. Minimum vector plotted: 2 cm s 1. meanders and eddies ( km). Southward branches off the AC are also evident at these longitudes with the strongest signature at about 25 W. East of 25 W, the circulation is dominated by a strong cyclonic cell centered in the GoC. Its formation in the model is caused by the forcing applied in the GoC and the use of the surface referenced potential density ( 0 ) as the vertical coordinate. A detailed examination of the cyclonic cell is given in sections 5c and 5d. b. The Azores Front at 30 W The model mean eastward velocity along 30 W is shown in Fig. 3. This section has been chosen for a detailed comparison of the model AF with observations and previous models. The AF studied by Gould (1985) based on a hydrographic survey is in the vicinity of this location. The distribution of the near-surface EKE and the potential vorticity field at this longitude are also available in the literature (e.g., Stammer and Böning 1996). The model AC is seen as a core of high speed eastward flow centered at about 33.5 N, extending down to 1000-m depth. The maximum mean speed reaches 9 cm s 1 near the surface. The eastward transport between 31 and 36 N is Sv, of which Sv is within the top 800 m. This value is in good agreement with several estimates of the AC transport from observations. Gould (1985) gives Sv for the upper 2000 m. An estimate of 10 Sv for the upper 800 m is given by Klein and Siedler (1989) at 35 W. Käse and Siedler (1982) obtained 10 Sv for the upper 1500 m southeast of the Azores near 22 W. The AC transport in the Canary Basin in the CME model (Spall 1990) is 6 7 Sv in the upper 800 m and 9 Sv for the upper 1500 m. Below 1000 m, there is a weak westward return flow underneath the model AC with maximum speed of about 1cms 1. North of the AC, there is a band of westward flow between 36 and 39 N with maximum speed in excess of 2 cm s 1 at about 500-m depth. Such a current is evident both in observations and in other high-resolution ocean models and has been termed the Azores Countercurrent by Onken (1993). It forms the southern limb of the anticyclonic circulation of the Subpolar Model Water (Pollard et al. 1996). The model EKE at 110-m depth and 30 W (averaged between 25 and 35 W) is shown in Fig. 4. Maximum variability is associated with the NAC ( 50 N) and the AC ( 33.5 N), in agreement with observed patterns from surface drifters and satellite altimetry. However, the absolute level of the EKE in the model is unrealistically low. For instance, the EKE for the model AC

6 SEPTEMBER 2000 JIA 2347 FIG. 2. The instantaneous velocity at 110 m in mid-july of the 20th year of the 1 3 model: (a) the GS region and (b) the Canary Basin. Minimum vector plotted: 2 cm s 1.

7 2348 JOURNAL OF PHYSICAL OCEANOGRAPHY FIG. 4. The meridional distribution of EKE at 110 m along 30 W (averaged between 25 and 35 W). FIG. 3. The mean eastward velocity at 30 W ofthe 1 3 model. Contour interval: 0.5 cm s 1. is 40 cm 2 s 2, while drifters give 200 cm 2 s 2 (Brügge 1995). For the NAC, the model EKE is an order of magnitude lower than that given by drifters. There is no clear explanation for this low energy level in the model. Possible factors include the lack of velocity shear in the mixed layer formulation, parameterization of subgrid-scale mixing and insufficient horizontal resolution. This model problem needs further investigation but will not be discussed here. The relatively high EKE level associated with the model AC is clearly seen in Fig. 4. It is significantly higher than the background level of variability and its magnitude is comparable with that of the model NAC. This pattern was not achieved by previous models at a similar resolution. The EKE in the CME models, for instance, has a pronounced maximum for the NAC, but is within the background level for the latitude of the AC (Beckmann et al. 1994a). Shown in Fig. 5 is the potential vorticity distribution at 30 W in the model. The front is marked by steeply sloping isopycnals at approximately 33.5 N. To the south of the front, there is a region of low potential vorticity at about 400 m, which corresponds to the subtropical mode water with potential density To the north of the front, there is also a region of low potential vorticity at about 600 m that corresponds to the subpolar mode water with potential density This pattern indicates that there is a reversal at depth of the meridional gradient of potential vorticity near 33.5 N, a favorable condition for baroclinic instability. 5. Further analysis and sensitivity experiments The presence of the AC and its associated meandering feature as well as the transport in the 1 3 model compare favorably with observations. The formation of the model AC is consistent with the proposed mechanism. It does not, however, exclude other mechanisms. To help provide further confirmation of the hypothesis, sensitivity experiments are performed at 4 3 resolution. They demonstrate the connection between the formation of the AC and the restoring condition in the GoC. The water mass transformation implied by the restoring condition in the GoC is examined and its ability to represent the Mediterranean overflow process is discussed. In addition, the mechanisms that determine the transport of the model AC and the causes of the strong cyclonic cell in the GoC are investigated. a. The Azores Current and the Mediterranean overflow Results from two experiments at 4 3 resolution are presented here to show the connection between the formation of the AC and the Mediterranean overflow. The first experiment (which will be identified as ACP for AC Present) is configured in a similar way to the 1 3 model (see section 3). It is not intended to provide a realistic simulation of the AC as the resolution is not sufficient to resolve the AC jet. It serves as a reference for the characteristics of the AC in the coarse resolution model, and provides a direct comparison with the 1 3 model. The second experiment (which will be identified as ACA for AC Absent) has the restoring condition removed from the GoC so that the effect of this forcing can be examined. The velocity field for experiment ACP at 110 m at the end of a 30-yr integration is shown in Fig. 6a. Instead of a jet structure, the model AC is represented by an enhanced eastward flow in the eastern basin between 30 and 40 N. There is no single origin identifiable in the GS for the model AC. The eastward flow is fed by

8 SEPTEMBER 2000 JIA 2349 FIG. 5. The mean potential vorticity and density distributions at 30 W (averaged between 25 and 35 W) in the 1 3 model. The potential vorticity is in gray scale in units of m 1 s 1. The density is contoured at model-layer densities. sources along the GS axis from 35 N, 70 W to45 N, 30 W, a pattern which is remarkably similar to the circulations on potential density surfaces and presented by Lozier et al. (1995). The mean position is in the southeast of the Grand Banks, which is the origin of the AC identified by Klein and Siedler (1989). In the GoC, the circulation is cyclonic but with a much weaker intensity than in the 1 3 model. South of the AC the flow is southward and forms part of the subtropical gyre recirculation. The eastward flow at 30 W (Fig. 7) is much slower and wider than in the 1 3 model but has a similar vertical extent. The AC transport at 30 W is 7.28 Sv evaluated between 30 and 40 N. Without the restoring condition in the GoC (see Fig. 6b for experiment ACA), the model AC is absent and the flow in the eastern basin is mostly southward. The southward flow is fed by branches off the NAC between approximately 40 and 50 N. As shown in Fig. 8, the accumulated barotropic transport (starting from the eastern boundary) at 25 and 36 N follows fairly well the Sverdrup relation based on a vertically integrated, steady, linear, inviscid vorticity balance. Away from the western boundary (east of 70 W), significant local departures occur only in regions where there are deep flows associated with topographic features. For example, in the model, the overflow across the Iceland Scotland Ridge runs southward along the eastern side of the Mid-Atlantic Ridge, this deep flow has caused the strong southward transport near 37 W at25 N and near 21 W at 36 N. The weak northward transport near 25 W at 25 N and near 18 W at36 N are caused by local recirculations associated with the deep southward flow. The effect of the restoring condition is small at 25 N, and away from the boundaries at 36 N. In the GoC, the restoring condition generates a strong cyclonic circulation near the eastern boundary (Fig. 8b). The cause of this barotropic circulation is investigated in section 5d. The two experiments (ACP and ACA) demonstrate that the formation of the model AC is a direct response to the restoring condition applied in the GoC. Away

9 2350 JOURNAL OF PHYSICAL OCEANOGRAPHY FIG. 7. The eastward velocity at 30 W at the end of a 30-yr integration of the 4 3 model experiment ACP. Contour interval: 0.2 cm s 1. FIG. 6. The velocity at 110 m at the end of a 30-yr integration of the 4 3 model: (a) experiment ACP with restoring in the GoC and (b) experiment ACA without restoring in the GoC. Minimum vector plotted: 1 cm s 1. from the restored region and the western boundary, the meridional flow is not significantly affected by the eastern boundary forcing. The next section investigates in detail the nature of the forcing implied by the restoring condition in the GoC. b. Water mass transformation in the Gulf of Cadiz The change in volume due to the restoring condition for each model layer is integrated for a period of one year for each of the three buffer zones (northern boundary, southern boundary, and the GoC). This volume is then converted to transport to give an estimate of the annual mean transport. If the net effect is to decrease the volume of a layer, it represents a sink to that layer. A sink at the eastern boundary (e.g., the GoC) may induce a net eastward transport (Pedlosky 1983, 1996). Similarly, a source at the eastern boundary implies a net westward transport. Shown in Fig. 9 is the transport in the GoC due to the restoring condition for the 1 3 model (Fig. 9a) and the 4 3 model experiment ACP (Fig. 9b). The transport patterns for both of the experiments are similar: eastward for densities to (model layers 7 to 14) and westward for density (model layer 15). The water mass transformation rate (from light to dense) is 4.26 Sv for the 1 3 model and 7.13 Sv for the 4 3 model. There are several estimates of the transport of Mediterranean Water near the Strait of Gibraltar and within the GoC after some mixing with the NACW. Ambar and Howe (1979) estimated an outflow transport of 1 Sv near the strait and 3 Sv near Cape St. Vincent. Ochoa and Bray (1991) estimated 0.5 Sv near the strait and 2.2 Sv near 8 W. More recently, Baringer and Price (1997) estimated that the Mediterranean outflow transport increases from about 0.7 Sv at the western end of the Strait of Gibraltar to about 1.9 Sv within the eastern GoC in the density range of , with a corresponding eastward transport of NACW in the density range Compared with the above observational estimates, the transformation rate in the 1 3 model is higher than the observations by about a factor of 2. It is even higher in the 4 3, by about a factor of 4. The larger value in the 4 3 model than in the 1 3 model is expected as the ability of the model to maintain the observed density structure is reduced because of the coarse resolution and the large diffusion. In density space, the water mass transformation in the

10 SEPTEMBER 2000 JIA 2351 FIG. 8. The accumulated barotropic transport (starting from the eastern boundary) for experiments ACA (thin solid line) and ACP (dashed line), and the Sverdrup transport derived from the ECMWF wind stress (thick solid line) (a) at 25 N and (b) at 36 N. models deviates significantly from observations. For example, the eastward transport for the 1 3 model in the GoC (Fig. 9a) is contained mostly within two density ranges: model layers 7 and 8 (densities and 27.03, combined transport 2.13 Sv) and model layers 12 to 14 (densities to 27.82, combined transport 2.13 Sv). Both the density and the transport in the shallower model layers correspond well to that of the NACW given by Baringer and Price (1997), but the presence of the eastward flow on the deeper layers is unrealistic. The restoring condition, however, does provide a westward transport for density (model layer 15), which is the layer just below the deepest restored layer interface, but it is too dense to be considered as the Mediterranean outflow water. The restoring condition, as implemented, tends to act as a sink for the layers above the deepest interface being restored and a source for the layer just below that. This peculiarity is caused by the use of surface-referenced potential density ( 0 ) in the model, which is also responsible for the presence of an over strong cyclonic cell in the GoC in the 1 3 model. This problem is discussed in more detail in section 5d. To partially correct the model deficiencies resulting from the restoring condition, a new experiment at 4 3 resolution (which will be identified as ACX) is performed with a simple modification. In this new experiment, the layer interface restoring applies only down to the upper interface of density layer (model layer 13), a shallower depth than previously. This is so that the mass input can be on a layer whose density (27.74) is within the range of the observed Mediterranean outflow water and that the eastward transport on the lower layers can be removed. The transport in the GoC due to the restoring condition for experiment ACX is shown in Fig. 9c. The westward transport is now mostly on density with a small contribution on density to give a total water mass transformation rate of 2.92 Sv, smaller in amplitude and shallower in depth than in experiment ACP. This reduction in the water mass transformation is largely due to the removal of the eastward transport on the deeper layers. The eastward transport of the NACW (densities to 27.22) is maintained with a small reduction (from 3.60 Sv in Fig. 9b to 2.92 Sv in Fig. 9c). This distribution is an improved representation of the observed pattern. The ideal situation would be for the mass input to be distributed among a range of densities (e.g., to 27.74) as presented by Baringer and Price (1997), but it is difficult to achieve with the present model formulation using surface-referenced potential density ( 0 ). The horizontal circulation at 110 m in experiment ACX (not shown here) is slightly weaker than that in experiment ACP in the western basin. In the eastern basin, the circulation patterns in both the experiments (ACP and ACX) are very similar. However, the vertical extent of the model AC at 30 W (Fig. 10) is shallower in experiment ACX in response to the changes in the eastern boundary and the AC transport is lower (3.53 Sv). c. The transport of the model Azores Current As shown in the last section, the restoring condition in the model induces a downward mass flux in the GoC. The oceanic response to this forcing in experiment ACP is an eastward flow into the boundary in the upper ocean (above 1700 m approximately) and a westward flow out of the boundary at depth (Fig. 11). This pattern is also true for experiment ACX except that the eastward flow extends to a shallower depth (1200 m approximately). This situation is consistent with the eastern boundary ventilation theory of Pedlosky (1983), an indication that the transport of the AC in the 4 3 model is essentially

11 2352 JOURNAL OF PHYSICAL OCEANOGRAPHY FIG. 9. The transport due to the restoring condition in the GoC in classes of model layer densities: (a) 1 3 model, (b) 4 3 model experiment ACP, and (c) 4 3 model experiment ACX. A positive value indicates a sink and corresponds to a net eastward transport. Note that only model layers 6 to 15 (density classes to 27.88) are plotted. Lighter layers have transport values close to zero as they outcrop to the mixed layer south of the GoC during most of the year cycle. The restoring condition does not apply to the mixed layer and the deeper layers. determined by the downward mass flux in the GoC (Table 3). A different circulation pattern is obtained in the 1 3 model (Fig. 12) where an additional horizontal recirculation is present both in the upper ocean (cyclonic) and at depth (anticyclonic). The cyclonic circulation penetrates down to about 1500 m and potential density 27.82, which covers the model layers with the sink from the restoring condition (model layers 7 to 14). The anticyclonic circulation below 1500 m is associated with the density (model layer 15), which has the source by the restoring condition. We also note that the strength of the flow at the southern part of the source and sink (between 33 and 35.5 N approximately) is stronger than the flow at the north (between 35.5 and 37 N approximately). Part of the eastward flow centered at about 33.5 N (the model AC) in the upper ocean feeds the westward return flow at depth. The rest is recircu- FIG. 10. The eastward velocity at 30 W at the end of a 30-yr integration of the 4 3 model experiment ACX. Contour interval: 0.2 cm s 1. FIG. 11. The baroclinic component of the zonal velocity at 10 W of the 4 3 model experiment ACP (taken from the snapshot at the end of a 30-yr integration). Contour interval: 0.4 cm s 1.

12 SEPTEMBER 2000 JIA 2353 TABLE 3. The deepest layer interface restored in the GoC, the AC transport at 30 W, and the water mass transformation rate in the GoC in the model experiments discussed in the text. The model experiment The deepest layer interface restored in the GoC The AC transport at 30 W (Sv) The water mass transformation rate in the GoC (Sv) 1 3 model ACP ACX lated horizontally with westward counterflow centered at about 36.5 N (which may be connected with the Azores Countercurrent downstream). It is the presence of this recirculation that enhances the transport of the AC. Such a circulation pattern is qualitatively consistent with the dipole circulation associated with isolated sources of Pedlosky (1996). We expect a weaker recirculation in the model than that predicted by the theory as the source and sink in the model is not isolated but situated against the eastern boundary. The theory predicts the upper limit of the transport of the zonal flows and an estimate is given below. Following Eq. (7.4.4) of Pedlosky (1996, p. 407), the transport of the eastward flow in the upper ocean (the AC) may be given as ym xw xw f y M 0 0 y S y x x TAC w dx dy w dx, S E E where w 0 is the vertical velocity (positive downward) estimated from the downward mass flux from the model upper layers to model layer 15 induced by the restoring forcing (w 0 is zero outside the restored area); f is the Coriolis parameter and is the meridional gradient of f; x E and x W are the longitudes bounding the restored area; y S is the latitude of the southern edge of the restored area; and y M is the latitude within the restored area where the eastward flow in the ocean upper layer reduces to zero or T AC reaches a maximum value, at approximately 35.5 N in Fig. 12. The first term in the equation is the contribution from the downward mass flux integrated over the area bounded by x E, x W, y S, and y M, it is estimated to be 2.47 Sv (if integrated over the entire restored area, it equals to the total water mass transformation of 4.26 Sv). The second term is the contribution from the horizontal recirculation (not the strength of the recirculation) and is the difference between the zonal integral evaluated at the two latitudes, y S and y M. Note that w 0 is zero at y S, so the second term in this case is the zonal integral at y M and is estimated to be Sv. Thus the eastward transport (T AC ) predicted by this theory is Sv, approximately 10 times of the total downward mass flux (4.26 Sv) in the GoC. The transport of the AC in the 1 3 model (10.50 Sv), although it exceeds the total mass transformation (4.26 FIG. 12. The baroclinic component of the zonal velocity at 10 W of the 1 3 model (taken from the instantaneous field in mid-july of the 20th year). Contour interval: 1 cm s 1. Sv), is much lower than the prediction, an indication that the presence of the boundary has a strong impact on the strength of the recirculation. The effect is even larger in the 4 3 model where the recirculation is absent. It is likely that the different resolutions mean that details of the restoring are different (e.g., the number of grid points within the restored area and along the coastline). Further investigations and experiments are necessary to understand the sensitivity of the zonal flows to the details of the boundary forcing, which will be the subject of future research. d. The cyclonic cell in the Gulf of Cadiz In the last section, it is shown that the downward mass flux induces a cyclonic circulation in the upper ocean in the GoC in the 1 3 model. This cyclonic circulation is much enhanced by a strong barotropic component as shown in Fig. 13. The vertically integrated transport in the cell reaches 20 Sv. This barotropic component is primarily caused by the use of surface-referenced potential density ( 0 ) as explained below. The use of 0 means that water masses at depth cannot be properly represented. This is particularly true for the Antarctic Bottom Water (AABW) in the deep ocean. Although it is denser than the North Atlantic Deep Water (NADW), which flows above, it has a smaller 0 value than the NADW due to the compressibility effect. As 0 is used as the vertical coordinate in the model, which needs to be monotonic with depth, the AABW and NADW of the initial ocean state need to be convectively adjusted. As a result, the signature of the AABW is missing in the model, and so is the northward transport associated with it. The transport at the model southern

13 2354 JOURNAL OF PHYSICAL OCEANOGRAPHY FIG. 13. The mean barotropic streamfunction (Sv) of the 1 3 model. Contour interval: 5 Sv. boundary due to the restoring condition is shown in Fig. 14 for the 1 3 model. There is northward transport of light water masses and southward transport of dense water masses, but the northward transport of AABW is missing. Such a pattern is also true for all the 4 3 model experiments presented earlier. FIG. 14. The transport due to the restoring condition at the southern boundary in classes of model layer densities. A positive value indicates a source (a net northward transport). The large southward transport in the bottom layers at the southern boundary has led to the thinning of these layers (in particular, densities and 27.92, model layers 15 and 16) during the course of the integration, which in turn, has resulted in the sinking of layer interfaces in the water column. This is particularly pronounced east of the Mid-Atlantic Ridge where there is a lack of mass input to these layers. West of the Mid- Atlantic Ridge, the export at the southern boundary is partly compensated by the overflows across the ridges connecting Greenland, Iceland, and Scotland. If a restoring condition is applied to layer interfaces in a small region, the interfaces will stay relatively high in the forced region and sink away from this region, thus a dome structure forms, which will drive a local cyclonic cell through geostrophy. Within the forced region, there will be an input of dense water to the bottom layers to compensate for the loss at the southern boundary in order to maintain the height of the layer interfaces. If the restoring condition only applies to the interfaces above a certain layer, the interfaces below will experience sinking, which will result in the thickening of the layer just below the deepest restored interface. This is precisely what happens in the 1 3 model in the GoC. The density distribution at 10 W is shown in Fig. 15 as an example. There are inversions of 0 at about 1000 m and below 2000 m in the Levitus climatology (Fig. 15a). These inversions are removed when the dataset is

14 SEPTEMBER 2000 JIA 2355 FIG. 16. The transport due to the restoring condition at the northern boundary in classes of model layer densities. A positive value indicates a sink (a net northward transport). used to initialize the model (Fig. 15b). After a 20-yr integration (Fig. 15c), the initial density distribution between and N and above the density contour at approximately 1500 m is reasonably maintained by the restoring condition. However, the sinking of the density contours to the north and south of the region and at depth is obvious. The dome structure is present and is most pronounced at intermediate depth, and acts to enhance the cyclonic circulation in the upper ocean in the GoC. The dome structure is much less pronounced in the 4 3 experiment ACP because the restoring is less effective in maintaining the observed density structure in the forced region and consequently generates a weaker barotropic cyclonic circulation in the GoC. Finally, for completeness, the transport at the northern boundary due to the restoring condition is shown in Fig. 16 for the 1 3 model. Light water masses are converted to dense water masses, which is consistent among all the experiments. It provides a reasonable representation of the water mass transformation occurring in the high latitudes of the North Atlantic. 6. Summary and discussion The principal issue addressed in this study is a mechanism for the formation of the AC. It is proposed that the AC can be induced by the water mass transformation associated with the Mediterranean overflow in the GoC. FIG. 15. The density distribution at 10 W: (a) Levitus climatology, (b) the initial state of the 1 3 model, and (c) at the end of a 20-yr integration of the 1 3 model. The density is contoured at model layer densities with two additional contours for and

15 2356 JOURNAL OF PHYSICAL OCEANOGRAPHY Observations show that the transport of Mediterranean outflow water through the Strait of Gibraltar increases significantly as it descends the continental slope by entraining the overlying NACW. This entrainment process creates a sink at the eastern boundary for the ocean upper layer in addition to the inflow into the Mediterranean. Existing theories (Luyten and Stommel 1986; Pedlosky 1983, 1996) suggest that a sink of this kind is capable of inducing zonal flows. This mechanism is confirmed by numerical experiments performed with and without the representation of the Mediterranean overflow process. The model does not include the Mediterranean overflow explicitly, but restores the model density fields in the GoC toward the observations. This restoring condition is shown to produce a reasonable representation of the water mass transformation deduced from observations. The AC transport is found to be comparable with the water mass transformation rate in the GoC in the coarse resolution model experiments. The eastward flowing AC into the boundary is returned as a westward flow at depth through the downward mass transfer imposed by the restoring condition in the GoC. This circulation pattern is consistent with the eastern boundary ventilation theory of Pedlosky (1983). At the eddy permitting resolution, the AC transport is larger than the water mass transformation rate in the GoC. This enhancement of the AC transport comes from the presence of an additional horizontal recirculation. Such a recirculation is predicted by the theory of dipole circulation associated with the isolated sources of Pedlosky (1996) although the strength of the circulation is much weaker in the model due to the presence of the boundary. This analysis suggests that the zonal flows are sensitive to the details of the forcing at the boundary. Other factors, though not investigated, may play a part in shaping the model AC. For example, the position of the GS separation may have influenced the separation of the GS into the NAC and the AC in the southeast of the Grand Banks. The model GS separates from the coast too far north due to the insufficient resolution to account for the complex western boundary processes (Özgökmen et al. 1997). It seems reasonable to speculate that the combined effect of western boundary processes (which govern the GS separation, its path downstream and its separation into the NAC and the AC) and the Mediterranean overflow process (which induces and maintains the AC and possibly encourages the separation of the GS into the NAC and the AC) largely determines the observed AC pathway. Thus it is essential to have these two processes properly represented before a realistic simulation of the AC can be achieved in general circulation models. Recent developments in high resolution modeling are promising in these respects. For instance, the high resolution simulation of the North Atlantic performed with MICOM at 1 12 resolution shows an improved GS separation and a realistic AC (E. Chassignet 1998, personal communication). This model uses a restoring condition in the GoC to represent the Mediterranean overflow process. Smith et al. (2000) also show an improved GS separation and a realistic AC in a North Atlantic model at 1 10 resolution using the Parallel Ocean Program model developed at Los Alamos National Laboratory. This model includes the Western Mediterranean Sea. Analysis of the results from these models should provide valuable information on the way in which the GS separation and the Mediterranean overflow affect the structure of the AC and its associated variability. The presence of the AC in response to the restoring condition imposed in the GoC suggests that the Mediterranean overflow is not only a source of warm and saline water at depth, but also has a strong dynamic impact on the upper-layer circulation in the subtropical eastern North Atlantic. A similar relationship also exists between the NAC and the overflows across the ridges connecting Greenland, Iceland, and Scotland. Böning et al. (1996) found that in the absence of the overflow across the Iceland Scotland Ridge the NAC develops a tendency to flow north toward Denmark Strait west of the Mid-Atlantic Ridge to be converted to Denmark Strait overflow water. This results in an unrealistic circulation pattern in the eastern North Atlantic. The path of the NAC improves with respect to its eastward penetration across the Mid-Atlantic Ridge when sufficient flow exchange is allowed across the Iceland Scotland Ridge (Semtner and Chervin 1992; Roberts and Wood 1997; Redler and Böning 1997). These studies demonstrate that overflows as a whole in the North Atlantic make a strong impact on the large-scale circulation in the upper ocean and emphasize that they need to be properly represented in ocean general circulation models. Further, overflows at high latitudes of the North Atlantic form an important part of the thermohaline circulation of the World Ocean (Broecker 1991), which may have significant effects on climate variability (Delworth et al. 1993). Acknowledgments. I am very grateful to Beverly de Cuevas and David Webb for their helpful comments on the draft manuscript. My sincere thanks go to the two anonymous reviewers who provided the most thorough and critical reviews of the manuscript. The 1 3 version of the model was developed and integrated under the DYNAMO project funded by the EU MAST Contract MAS2-CT The contributions from Sally Barnard and Adrian New to the project are much appreciated. REFERENCES Ambar, I., and M. R. Howe, 1979: Observations of the Mediterranean outflow. II. The deep circulation in the vicinity of the Gulf of Cadiz. Deep-Sea Res., 26A, Baringer, M. O., and J. F. Price, 1997: Mixing and spreading of the Mediterranean outflow. J. Phys. Oceanogr., 27, Barnier, B., L. Siefridt, and P. Marchesiello, 1995: Thermal forcing