Formation of an Azores Current Due to Mediterranean Overflow in A

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2342 JOURNAL OF PHYSICAL OCEANOGRAPHY VOLUME 30

Formation of an Azores Current Due to Mediterranean Over¯ow in a Modeling Study of the North Atlantic

YANLI JIA Southampton Oceanography Centre, Southampton, United Kingdom

(Manuscript received 24 August 1998, in ®nal 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 over¯ow in the Gulf of Cadiz. Observations show that the transport of the Mediterranean out¯ow water through the Strait of Gibraltar increases signi®cantly 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 in¯ow into the Mediterranean. Such a sink is capable of inducing strong zonal ¯ows such as the Azores Current. This mechanism is con®rmed by numerical experiments with and without the representation of the Mediterranean over¯ow process. The numerical model is based on the Miami Isopycnic Coordinate Ocean Model. The model does not include the Mediterranean over¯ow explicitly, but restores the model density ®elds 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 over¯ow 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 Medi- terranean over¯ow process in general circulation models in order to capture the correct characteristics of the ¯ow ®elds and water masses in the subtropical eastern North Atlantic.

1. Introduction the African coast with southward branches in the Canary Basin as part of the subtropical gyre recirculation Southeast of the Grand Banks of Newfoundland the (Stramma 1984; Olbers et al. 1985; Klein and Siedler Gulf Stream (GS) separates into two branches. The 1989). The hydrographic database of Lozier et al. (1995) northern branch turns northeastward and becomes the reveals a coherent AC that stretches across the eastern North Atlantic Current (NAC). The southern branch, which becomes the Azores Current (AC), heads south- half of the basin, with divergences to the south and eastward across the Mid-Atlantic Ridge to the south of convergences from the north such that the downstream the Azores, then ¯ows mainly eastward at a latitude of transport does not change much. Recent hydrographic about 35ЊN to the Gulf of Cadiz (GoC). surveys also indicate the eastward extension of the AC Associated with the AC is a front with signi®cant to the Moroccan continental slopes (FernaÂndez and Pin- temperature and salinity contrasts. There have been a gree 1996; Pingree 1997). Buoys deployed in the AC number of detailed hydrographic surveys of the front at are found to travel eastward and reach the western side various locations, for example, to the southeast of the of the GoC, and then move northward or southward Grand Banks (Mann 1967, 1972; Clarke et al. 1980), along the continental slopes. in the region of the Mid-Atlantic Ridge (Gould 1985; Based on the above surveys, the AC is observed to Sy 1988; Stramma and Muller 1989), and southeast of be a meandering jet 60±100 km wide with an eastward the Azores (KaÈse and Siedler 1982; KaÈse et al. 1985; velocity of 25±50 cm sϪ1. The eastward ¯ow is mostly Siedler et al. 1985; Rios et al. 1992). Geostrophic trans- in the upper few hundred meters but can reach as deep port ®elds obtained from historical hydrographic data as 2000 m. The current carries a large fraction of the indicate that the eastward ¯ow extends all the way to water entering the eastern recirculation region of the Canary Basin. The estimates of the AC transport are in the range of 10±15 Sv (Sv ϵ 106 m3 sϪ1). The surface temperature and salinity changes across the front can Corresponding author address: Dr. Yanli Jia, Southampton Ocean- be as large as 2 C and 0.3 psu. The front marks the ography Centre, Empress Dock, Southampton SO14 3ZH, United Њ Kingdom. northern boundary of the 18ЊC Sargasso Sea water in E-mail: [email protected] the central North Atlantic.

᭧ 2000 American Meteorological Society

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Both drifter data (Richardson 1983; Krauss and KaÈse argued that under such a variable condition, nonlinear 1994; BruÈgge 1995) and satellite altimetry (e.g., Le inertial effects may become the dominant dynamic con- Traon et al. 1990; Wunsch and Stammer 1995; Stammer trol in the GS extension region, thus the separation of 1997) show a band of high eddy kinetic energy (EKE) the GS into the NAC and the AC. They also suggested associated with the AC. KaÈse and Siedler (1982) ob- that a local relative minimum in the magnitude of the served considerable meandering of the front southeast wind stress curl exists in the inner subtropical region, of the Azores with mesoscale eddies on both sides of which permits a quasi-zonal ¯ow to continue to the the front. Baroclinic instability has been identi®ed as eastern basin, hence the eastward extension of the AC. one of the mechanisms for the high energy level (Beck- In this study, a complementary mechanism for the mann et al. 1994a). The isopycnic potential vorticity formation and maintenance of the AC is proposed. It is analysis of climatological hydrographic data (Stammer suggested that the water mass transformation associated and Woods 1987) indicates that a necessary condition with the Mediterranean over¯ow in the GoC may induce for baroclinic instability, the reversal at depth of the the AC. The streamfunction ®eld for potential density meridional gradient of potential vorticity, is present at surface ␴ 0 ϭ 27.00 in Lozier et al. (1995) clearly shows the AC. the convergence of the streamlines associated with the The primary mechanism for the formation and main- AC in the GoC, which suggests a possible connection tenance of the Azores Front (AF) is unknown. For this between the two. The Mediterranean over¯ow is often very reason, numerical modeling of the AF and its as- poorly represented in general circulation models, which sociated variability has been dif®cult. The front is either may be the cause for a nonexistent AC in many cases. absent or very weak in general circulation models. For The AC in the eastern basin and the GoC occupy a instance, the high resolution model of the North Atlantic similar latitudinal extent and this may not be coinci- developed under the Community Modelling Effort dental. The following is an argument to explain how (CME) does not produce the separation of the GS into this may operate. the NAC and the AC (Bryan and Holland 1989). The At the Strait of Gibraltar dense Mediterranean water model AC develops only in the eastern basin (Spall spills over the sills into the North Atlantic. The transport 1990). The EKE at the AC latitudes is barely above the of the Mediterranean out¯ow water at the western end background level of variability (Treguier 1992). The of the Strait of Gibraltar is typically 1 Sv (Lacombe and meridional density gradient associated with the model Richez 1982; Bryden et al. 1994; Baringer and Price AC is too weak and the necessary condition for baro- 1997). Intense mixing in the GoC increases this trans- clinic instability is not satis®ed (Beckmann et al. 1994a). port by a factor of about 3 by entraining the overlying No signi®cant improvement is found with increased hor- North Atlantic Central Water (NACW) (Ambar and izontal resolution (Beckmann et al. 1994b). Howe 1979; Ochoa and Bray 1991; Baringer and Price This paper reports the occurrence of an AC in a gen- 1997). This entrainment process in the GoC introduces eral circulation model of the North Atlantic. A mech- a sink at the eastern boundary to the ocean upper layer anism for the formation of the AC is ®rst proposed. The in addition to the in¯ow into the Mediterranean. Such characteristics of the AC in the model are then presented a sink is capable of inducing strong zonal ¯ows such and compared with what we know from available ob- as the AC. servations. Further sensitivity experiments are per- There have been earlier laboratory experiments and formed to verify the proposed mechanism. 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 2. A mechanism for the formation of the (extraction) of ¯ow into (out of) the system, or through Azores Current vertical ¯ux of mass across density surfaces. In a lab- The presence of the AC, a coherent zonal ¯ow in the oratory experiment, Stommel et al. (1958) showed the inner subtropical gyre that stretches across a large extent induction of basin-scale zonal ¯ows by a point source of the basin, is a major feature of the circulation of the and sink placed near the eastern boundary in a rotating North Atlantic. It is situated well to the south of the system. In a theoretical analysis, Pedlosky (1996) mean zero wind stress curl where Ekman pumping im- showed that, for a localized source or sink of ®nite plies southward transport in the ocean, thus its zonal extent situated some distance away from the lateral orientation cannot be fully explained by Sverdrup dy- boundaries, zonal ¯ows form west of and within the namics. There must be other mechanisms involved. One latitudinal band of the source or sink. The zonal ¯ows interpretation was given by KaÈse and Krauss (1996) must be bidirectional under the constraint of zero pres- through a close examination of a time series of the wind sure gradient east of the source or sink and outside the stress curl over the North Atlantic (monthly mean and latitudinal band containing the source or sink. The trans- zonally averaged between 65Њ and 5ЊW). They showed port of each of the zonal ¯ows is large in relation to that in the latitudinal band between 35Њ and 50ЊN, the the strength of the source or sink, but the difference in wind stress curl, though signi®cantly negative, is highly the zonal ¯ows is equal to the source or sink. variable with annual and interannual ¯uctuations. They A similar circulation pattern was obtained by Luyten

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Ϫ1 TABLE 1. The surface-referenced potential density anomaly (␴0)of TABLE 2. The diffusion velocities (cm s ) for momentum, tracers 1 4 the 20 layers de®ning the discretised vertical coordinate of the model. and layer thickness for the ⁄3Њ and the ⁄3Њ models. Each diffusion Layer 1 is the model mixed layer that has a variable density distri- coef®cient is the product of the appropriate diffusion velocity and bution. the grid spacing.

Layer ␴0 Layer ␴0 Momentum Thickness Tracers

1 1 Variable 11 27.52 ⁄3Њ model 0.5 0.1 0.5 4 2 24.70 12 27.64 ⁄3Њ model 1.0 0.5 0.5 3 25.28 13 27.74 4 25.77 14 27.82 5 26.18 15 27.88 6 26.52 16 27.92 device to simulate the observed transport pattern as- 7 26.80 17 28.00 sociated with the Mediterranean over¯ow. Sensitivity 8 27.03 18 28.06 experiments show that this restoring condition is the 9 27.22 19 28.09 10 27.38 20 28.12 driving force for the model AC. When this condition is removed, no AC is formed.

3. Model description and Stommel (1986) in a two moving layer, two gyre ocean circulation model where a downward interfacial The model used in this study is based on the Miami (buoyancy) ¯ux was applied within a circle of ®nite Isopycnic Coordinate Ocean Model (MICOM) using radius in the eastern basin but away from the eastern surface-referenced potential density (␴ 0) as the vertical boundary. Dipole circulation forms on both the upper coordinate. The details of the model numerics can be and lower moving layers west of the circle and merid- found in Bleck et al. (1992). The model domain covers ional ¯ows exist only inside the circle. A different ¯ow the North Atlantic basin from approximately 20ЊSto pattern is predicted by the eastern boundary ventilation 70ЊN and from 100ЊWto20ЊE. The horizontal resolution 1 1 theory of Pedlosky (1983). By allowing a downward is ⁄3Њ in longitude by ⁄3Њ cos(␾) in latitude (where ␾ is transfer of mass at the eastern boundary, an eastward latitude), thus yielding an isotropic horizontal grid. ¯ow into the boundary results in the upper layer and a There are 20 layers in the vertical. The top layer is a westward ¯ow out of the boundary in the lower layer. mixed layer of Kraus±Turner formulation where density The ¯ow is unidirectional within each layer and the and other model variables are allowed to vary with time transport matches the downward mass transfer. and in space. The 19 layers below the mixed layer are The entrainment process by the sinking of the Med- layers with constant potential density. The values of the iterranean over¯ow in the GoC involves a downward layer densities (Table 1) are chosen so that the water mass transfer in a region attached to the eastern bound- masses and the associated dynamics in the North At- ary. The source (in the lower layer) and the sink (in the lantic can be represented as well as possible with the upper layer) are ®nite but not isolated. We may envisage limited resolution. The bathymetry is taken from the a situation where both the above forcing mechanisms ETOPO5 database from the National Geophysical Data (Luyten and Stommel 1986; Pedlosky 1983) operate in Centre; no smoothing is applied but a minimum ocean the system and set the upper and lower limits on the depth of 75 m is set in the model. strength of the zonal ¯ows west of the GoC. Subgrid-scale processes are parameterized in Lapla- The ¯ow exchange at the Strait of Gibraltar may in- cian form. There are three lateral mixing parameters for troduce an additional forcing to the zonal ¯ows. The isopycnic diffusion of momentum, layer thickness, and theoretical model of Webb (1993), which considers a tracers. They are written in the form of diffusion ve- layer of ¯uid in the ocean interior including horizontal locities and represent the ratio of the diffusion coef®- viscosity but neglecting vertical viscosity, suggests that cients to the model horizontal grid spacing. These dif- a point source or sink at the eastern boundary results fusion velocities are kept constant in the model (Table in an east±west jet, which becomes wider toward the 2), thus diffusion is proportional to grid spacing. The west. Within the jet, there is a balance between viscosity, diapycnic mixing coef®cient for tracers is a function of which works to spread the jet, and Rossby waves prop- strati®cation. It is inversely proportional to the Brunt± Ϫ7 2 agating from the eastern boundary, which works to con- VaÈisaÈlaÈ frequency (N): a 0/N, where a 0 is set to 10 m strain it in a north±south direction. The transport of the sϪ2. There is no diapycnic mixing of momentum be- jet is determined by the strength of the source or sink. tween layers. In the present modeling study, the Mediterranean The wind stress, friction velocity (used in computing over¯ow process is not explicitly resolved but is rep- turbulent kinetic energy for mixed layer forcing) and resented by restoring the model density ®eld toward the heat ¯ux used to force the model are taken from a observed values in the GoC. It is shown that this re- three-year monthly climatology derived from the storing condition is capable of extracting lighter water ECMWF analysis (1986±88). The full surface heat ¯ux from the upper ocean and replacing it with heavier water consists of the ECMWF climatology plus a restoring at depth, thus serving as a water mass transformation term toward an equivalent sea surface temperature with

Unauthenticated | Downloaded 09/24/21 02:10 PM UTC SEPTEMBER 2000 JIA 2345 a variable timescale as described in Barnier et al. (1995). of 20 years. The model ®elds are saved every three days The surface salinity is restored toward the Levitus for the last ®ve years of the integration and are used to (1982) climatology with a timescale identical to that calculate the time mean ®elds. 4 used in the heat ¯ux formulation. A coarse resolution ( ⁄3Њ) version of the model is set The model boundaries are closed in the north, south, up in a similar way. The mixing parameters are given and at the Strait of Gibraltar. Connection with the ocean in Table 2. The sizes of the buffer zones are the same 1 exterior to the model domain is parameterized by means as in the ⁄3Њ model. This model is used for sensitivity of buffer zones in which model variables (depth of in- experiments presented in section 5. 1 terfaces below the base of the mixed layer and salinity The high resolution ( ⁄3Њ) version of the model was of the density layers) are restored toward observed val- developed and integrated under the DYNAMO (Dy- ues. The restoring of layer interface depth is designed namics of North Atlantic Models) project funded by the to account for the processes that determine the observed European Union. It was used in an extensive modeling density structure in the buffer zones. The inclusion of intercomparison exercise involving three models. A de- salinity restoring on the density layers does not in¯uence tailed account of the model intercomparison can be the dynamics (as the density of each layer is ®xed), but found in DYNAMO Group (1997). provides the layers with the appropriate water mass characteristics (temperature and salinity). In the version Model results: 1 ؇ model .4 of MICOM used in this study, salinity is the active ⁄3 thermodynamic variable in the density layers, which is In this section, the characteristics of the AC in the 1 advected and diffused. Layer temperature is derived ⁄3Њ model experiment are presented and compared with from layer density and salinity. observations. Both instantaneous and mean ®elds are The northern buffer zone covers the region north of shown. Unless stated otherwise, the mean ®elds are the 67ЊN between 40Њ and 10ЊW, and north of the line con- annual averages over the last ®ve years of the model necting 67ЊN, 10ЊW and 60ЊN, 17.5ЊE (17.5ЊE is the integration. model eastern boundary). Data used for the restoring in this region is based on a hydrographic database from a. Near-surface circulation the National Oceanographic Data Center (NODC In- formal Report 12, 1991). This database provides an im- The mean circulation pattern at 110 m is depicted in proved description of the water mass characteristics in Fig. 1. The major currents of the GS system are all this region. The combination of a relatively large buffer present: the Florida Current, the GS, the NAC, and the zone and an improved hydrographic database is de- AC. The model shows a signi®cant improvement over signed to achieve a realistic representation of water mass previous models in its representation of the AC. How- transformation occurring north of the ridges connecting ever, the GS separates from the coast too far north Greenland, Iceland, and Scotland. (ϳ40ЊN) compared with observations (ϳ35ЊN). Con- The Levitus (1982) climatology is used for the re- sequently, downstream, the separation of the GS into storing in the southern boundary buffer zone and in the the NAC and the AC is unrealistic. There are two iden- GoC. The southern buffer zone is de®ned south of ti®able origins for the AC, one at approximately 42ЊN, 11.5ЊS. The Strait of Gibraltar in the model is at 36ЊN, 47ЊW southeast of the Grand Banks and the other at 6ЊW and the associated buffer zone is bounded by Cape Hatteras at about 35ЊN. The latter is not present 33.5ЊN and 38.0ЊN, 11ЊW and 6ЊW. in observations although it appears to be the major Full depth restoring is applied in the northern and source for the model AC. The former is in agreement southern buffer zones. The restoring timescale increases with the analysis by Klein and Siedler (1989) based on linearly from 3 days at the model boundaries (70ЊN and historical hydrographic data, although the precise po- 19ЊS) to 100 days at the inner edges of the buffer zones. sition is slightly too far north. According to Klein and In the GoC, the restoring condition applies down to the Siedler (1989), there is a direct current connecting the upper interface of density layer 27.88 (model layer 15) source region (approximately 40ЊN, 45ЊW) and the AC and to the salinity of the layers above that interface (at region southwest of the Azores in winter, while in sum- approximately 1500-m depth). The timescale increases mer the ¯ow separates from the source region into two linearly from 14 days at the Strait of Gibraltar to 100 branches, one ¯owing directly towards the AC region days at a distance of 300 km. and the other taking a cyclonic loop before joining the Levitus (1982) September climatology is used to ini- AC. This loop is present in the model throughout the tialize the model except for the region north of 60ЊN year with slight strengthening in summer. The direct and east of 40ЊW where, for consistency, the same da- route from the origin to the AC is missing from the tabase as for the northern boundary restoring is used. model. Initially the North Atlantic basin is mostly occupied by The model AC appears as a wide zonal jet in the the top 16 layers; the bottom 4 layers reside mainly mean ®eld between 32Њ and 35ЊN, 50Њ and 25ЊW. The north of the ridges connecting Greenland, Iceland, and instantaneous ®eld (Fig. 2) exhibits a much more vig- Scotland. The model is then integrated for a total length orous ¯ow pattern with a tighter jet (150 km) and large

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1 Ϫ1 FIG. 1. The mean velocity at 110 m of the ⁄3Њ model. Minimum vector plotted: 2 cm s . meanders and eddies (500±600 km). Southward branch- the top 800 m. This value is in good agreement with es off the AC are also evident at these longitudes with several estimates of the AC transport from observations. the strongest signature at about 25ЊW. Gould (1985) gives 10±12 Sv for the upper 2000 m. East of 25ЊW, the circulation is dominated by a strong An estimate of 10 Sv for the upper 800 m is given by cyclonic cell centered in the GoC. Its formation in the Klein and Siedler (1989) at 35ЊW. KaÈse and Siedler model is caused by the forcing applied in the GoC and (1982) obtained 10 Sv for the upper 1500 m southeast the use of the surface referenced potential density (␴ 0) of the Azores near 22ЊW. The AC transport in the Canary as the vertical coordinate. A detailed examination of the Basin in the CME model (Spall 1990) is 6±7 Sv in the cyclonic cell is given in sections 5c and 5d. upper 800 m and 9 Sv for the upper 1500 m. Below 1000 m, there is a weak westward return ¯ow underneath the model AC with maximum speed of about b. The Azores Front at 30ЊW 1cmsϪ1. North of the AC, there is a band of westward The model mean eastward velocity along 30ЊWis ¯ow between 36Њ and 39ЊN with maximum speed in shown in Fig. 3. This section has been chosen for a excess of 2 cm sϪ1 at about 500-m depth. Such a current detailed comparison of the model AF with observations is evident both in observations and in other high-res- and previous models. The AF studied by Gould (1985) olution ocean models and has been termed the Azores based on a hydrographic survey is in the vicinity of this Countercurrent by Onken (1993). It forms the southern location. The distribution of the near-surface EKE and limb of the anticyclonic circulation of the ``Subpolar the potential vorticity ®eld at this longitude are also Model Water'' (Pollard et al. 1996). available in the literature (e.g., Stammer and BoÈning The model EKE at 110-m depth and 30ЊW (averaged 1996). between 25Њ and 35ЊW) is shown in Fig. 4. Maximum The model AC is seen as a core of high speed east- variability is associated with the NAC (ϳ50ЊN) and the ward ¯ow centered at about 33.5ЊN, extending down to AC (ϳ33.5ЊN), in agreement with observed patterns 1000-m depth. The maximum mean speed reaches 9 cm from surface drifters and satellite altimetry. However, sϪ1 near the surface. The eastward transport between the absolute level of the EKE in the model is unreal- 31Њ and 36ЊN is 10.50 Sv, of which 10.00 Sv is within istically low. For instance, the EKE for the model AC

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1 FIG. 2. The instantaneous velocity at 110 m in mid-July of the 20th year of the ⁄3Њ model: (a) the GS region and (b) the Canary Basin. Minimum vector plotted: 2 cm sϪ1.

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FIG. 4. The meridional distribution of EKE at 110 m along 30ЊW (averaged between 25Њ and 35ЊW).

1 FIG. 3. The mean eastward velocity at 30ЊWofthe ⁄3Њ model. Con- tour interval: 0.5 cm sϪ1. favorably with observations. The formation of the model AC is consistent with the proposed mechanism. It is 40 cm2 sϪ2, while drifters give 200 cm2 sϪ2 (BruÈgge does not, however, exclude other mechanisms. To help 1995). For the NAC, the model EKE is an order of provide further con®rmation of the hypothesis, sensi- 4 magnitude lower than that given by drifters. There is tivity experiments are performed at ⁄3Њ resolution. They no clear explanation for this low energy level in the demonstrate the connection between the formation of model. Possible factors include the lack of velocity shear the AC and the restoring condition in the GoC. The in the mixed layer formulation, parameterization of sub- water mass transformation implied by the restoring con- grid-scale mixing and insuf®cient horizontal resolution. dition in the GoC is examined and its ability to represent This model problem needs further investigation but will the Mediterranean over¯ow process is discussed. In ad- not be discussed here. dition, the mechanisms that determine the transport of The relatively high EKE level associated with the the model AC and the causes of the strong cyclonic cell model AC is clearly seen in Fig. 4. It is signi®cantly in the GoC are investigated. higher than the background level of variability and its magnitude is comparable with that of the model NAC. a. The Azores Current and the Mediterranean This pattern was not achieved by previous models at a over¯ow similar resolution. The EKE in the CME models, for 4 instance, has a pronounced maximum for the NAC, but Results from two experiments at ⁄3Њ resolution are is within the background level for the latitude of the presented here to show the connection between the for- AC (Beckmann et al. 1994a). Shown in Fig. 5 is the mation of the AC and the Mediterranean over¯ow. The potential vorticity distribution at 30ЊW in the model. ®rst experiment (which will be identi®ed as ACP for 1 The front is marked by steeply sloping isopycnals at AC Present) is con®gured in a similar way to the ⁄3Њ approximately 33.5ЊN. To the south of the front, there model (see section 3). It is not intended to provide a is a region of low potential vorticity at about 400 m, realistic simulation of the AC as the resolution is not which corresponds to the subtropical mode water with suf®cient to resolve the AC jet. It serves as a reference potential density 26.80. To the north of the front, there for the characteristics of the AC in the coarse resolution 1 is also a region of low potential vorticity at about 600 model, and provides a direct comparison with the ⁄3Њ m that corresponds to the subpolar mode water with model. The second experiment (which will be identi®ed potential density 27.22. This pattern indicates that there as ACA for AC Absent) has the restoring condition is a reversal at depth of the meridional gradient of po- removed from the GoC so that the effect of this forcing tential vorticity near 33.5ЊN, a favorable condition for can be examined. baroclinic instability. The velocity ®eld 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 5. Further analysis and sensitivity experiments enhanced eastward ¯ow in the eastern basin between The presence of the AC and its associated meandering 30Њ and 40ЊN. There is no single origin identi®able in 1 feature as well as the transport in the ⁄3Њ model compare the GS for the model AC. The eastward ¯ow is fed by

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FIG. 5. The mean potential vorticity and density distributions at 30ЊW (averaged between 25Њ and 1 Ϫ11 Ϫ1 Ϫ1 35ЊW) in the ⁄3Њ model. The potential vorticity is in gray scale in units of 10 m s . The density is contoured at model-layer densities. sources along the GS axis from 35ЊN, 70ЊWto45ЊN, Sverdrup relation based on a vertically integrated, 30ЊW, a pattern which is remarkably similar to the cir- steady, linear, inviscid vorticity balance. Away from the culations on potential density surfaces ␴ 0 ϭ 26.50 and western boundary (east of 70ЊW), signi®cant local de- ␴ 0 ϭ 27.00 presented by Lozier et al. (1995). The mean partures occur only in regions where there are deep position is in the southeast of the Grand Banks, which ¯ows associated with topographic features. For exam- is the origin of the AC identi®ed by Klein and Siedler ple, in the model, the over¯ow across the Iceland±Scot- (1989). In the GoC, the circulation is cyclonic but with land Ridge runs southward along the eastern side of the 1 a much weaker intensity than in the ⁄3Њ model. South Mid-Atlantic Ridge, this deep ¯ow has caused the strong of the AC the ¯ow is southward and forms part of the southward transport near 37ЊWat25ЊN and near 21ЊW subtropical gyre recirculation. The eastward ¯ow at at 36ЊN. The weak northward transport near 25ЊWat 30ЊW (Fig. 7) is much slower and wider than in the 25ЊN and near 18ЊWat36ЊN are caused by local re- 1 ⁄3Њ model but has a similar vertical extent. The AC circulations associated with the deep southward ¯ow. transport at 30ЊW is 7.28 Sv evaluated between 30Њ and The effect of the restoring condition is small at 25ЊN, 40ЊN. and away from the boundaries at 36ЊN. In the GoC, the Without the restoring condition in the GoC (see Fig. restoring condition generates a strong cyclonic circu- 6b for experiment ACA), the model AC is absent and lation near the eastern boundary (Fig. 8b). The cause the ¯ow in the eastern basin is mostly southward. The of this barotropic circulation is investigated in section southward ¯ow is fed by branches off the NAC between 5d. approximately 40Њ and 50ЊN. As shown in Fig. 8, the The two experiments (ACP and ACA) demonstrate accumulated barotropic transport (starting from the east- that the formation of the model AC is a direct response ern boundary) at 25Њ and 36ЊN follows fairly well the to the restoring condition applied in the GoC. Away

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FIG. 7. The eastward velocity at 30ЊW at the end of a 30-yr in- 4 tegration of the ⁄3Њ model experiment ACP. Contour interval: 0.2 cm sϪ1.

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 1 in the GoC due to the restoring condition for the ⁄3Њ 4 model (Fig. 9a) and the ⁄3Њ model experiment ACP (Fig. 9b). The transport patterns for both of the experiments are similar: eastward for densities 26.80 to 27.82 (model layers 7 to 14) and westward for density 27.88 (model layer 15). The water mass transformation rate (from 1 light to dense) is 4.26 Sv for the ⁄3Њ model and 7.13 Sv 4 for the ⁄3Њ model. There are several estimates of the transport of Med- iterranean Water near the Strait of Gibraltar and within FIG. 6. The velocity at 110 m at the end of a 30-yr integration of 4 the GoC after some mixing with the NACW. Ambar and the ⁄3Њ model: (a) experiment ACP with restoring in the GoC and (b) experiment ACA without restoring in the GoC. Minimum vector plot- Howe (1979) estimated an out¯ow transport of 1 Sv ted: 1 cm sϪ1. 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 8ЊW. More recently, Baringer and Price from the restored region and the western boundary, the (1997) estimated that the Mediterranean out¯ow trans- meridional ¯ow is not signi®cantly affected by the east- port increases from about 0.7 Sv at the western end of ern boundary forcing. The next section investigates in the Strait of Gibraltar to about 1.9 Sv within the eastern detail the nature of the forcing implied by the restoring GoC in the density range of 27.3±28.0, with a corre- condition in the GoC. sponding eastward transport of NACW in the density range 26.6±27.3. Compared with the above observational estimates, the b. Water mass transformation in the Gulf of Cadiz 1 transformation rate in the ⁄3Њ model is higher than the The change in volume due to the restoring condition observations by about a factor of 2. It is even higher in 4 for each model layer is integrated for a period of one the ⁄3Њ, by about a factor of 4. The larger value in the 4 1 year for each of the three buffer zones (northern bound- ⁄3Њ model than in the ⁄3Њ model is expected as the ability ary, southern boundary, and the GoC). This volume is of the model to maintain the observed density structure then converted to transport to give an estimate of the is reduced because of the coarse resolution and the large annual mean transport. If the net effect is to decrease diffusion. the volume of a layer, it represents a sink to that layer. In density space, the water mass transformation in the

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peculiarity is caused by the use of surface-referenced

potential density (␴ 0) in the model, which is also re- sponsible for the presence of an over strong cyclonic 1 cell in the GoC in the ⁄3Њ model. This problem is discussed in more detail in section 5d. To partially correct the model de®ciencies resulting 4 from the restoring condition, a new experiment at ⁄3Њ resolution (which will be identi®ed as ACX) is performed with a simple modi®cation. In this new experiment, the layer interface restoring applies only down to the upper interface of density layer 27.74 (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 out- ¯ow 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 27.74 with a small contribution on density 27.64 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 26.80 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., 27.38 to 27.74) as presented by Baringer and Price (1997), but it is dif®cult to achieve with the present model formulation using surface-referenced po-

tential density (␴ 0). FIG. 8. The accumulated barotropic transport (starting from the The horizontal circulation at 110 m in experiment eastern boundary) for experiments ACA (thin solid line) and ACP ACX (not shown here) is slightly weaker than that in (dashed line), and the Sverdrup transport derived from the ECMWF experiment ACP in the western basin. In the eastern wind stress (thick solid line) (a) at 25ЊN and (b) at 36ЊN. basin, the circulation patterns in both the experiments (ACP and ACX) are very similar. However, the vertical extent of the model AC at 30ЊW (Fig. 10) is shallower models deviates signi®cantly from observations. For ex- in experiment ACX in response to the changes in the 1 ample, the eastward transport for the ⁄3Њ model in the eastern boundary and the AC transport is lower (3.53 GoC (Fig. 9a) is contained mostly within two density Sv). ranges: model layers 7 and 8 (densities 26.80 and 27.03, combined transport 2.13 Sv) and model layers 12 to 14 (densities 27.64 to 27.82, combined transport 2.13 Sv). c. The transport of the model Azores Current Both the density and the transport in the shallower model layers correspond well to that of the NACW given As shown in the last section, the restoring condition by Baringer and Price (1997), but the presence of the in the model induces a downward mass ¯ux in the GoC. eastward ¯ow on the deeper layers is unrealistic. The The oceanic response to this forcing in experiment ACP restoring condition, however, does provide a westward is an eastward ¯ow into the boundary in the upper ocean transport for density 27.88 (model layer 15), which is (above 1700 m approximately) and a westward ¯ow out the layer just below the deepest restored layer interface, of the boundary at depth (Fig. 11). This pattern is also but it is too dense to be considered as the Mediterranean true for experiment ACX except that the eastward ¯ow out¯ow water. extends to a shallower depth (1200 m approximately). The restoring condition, as implemented, tends to act This situation is consistent with the eastern boundary as a sink for the layers above the deepest interface being ventilation theory of Pedlosky (1983), an indication that 4 restored and a source for the layer just below that. This the transport of the AC in the ⁄3Њ model is essentially

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FIG. 9. The transport due to the restoring condition in the GoC in classes of model layer densities: (a) 1 4 4 ⁄3Њ model, (b) ⁄3Њ model experiment ACP, and (c) ⁄3Њ 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 26.52 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 ¯ux in the GoC (Ta- ticyclonic circulation below 1500 m is associated with ble 3). the density 27.88 (model layer 15), which has the source 1 A different circulation pattern is obtained in the ⁄3Њ by the restoring condition. We also note that the strength model (Fig. 12) where an additional horizontal recir- of the ¯ow at the southern part of the source and sink culation is present both in the upper ocean (cyclonic) (between 33Њ and 35.5ЊN approximately) is stronger and at depth (anticyclonic). The cyclonic circulation than the ¯ow at the north (between 35.5Њ and 37ЊN penetrates down to about 1500 m and potential density approximately). Part of the eastward ¯ow centered at 27.82, which covers the model layers with the sink from about 33.5ЊN (the model AC) in the upper ocean feeds the restoring condition (model layers 7 to 14). The an- the westward return ¯ow at depth. The rest is recircu-

FIG. 10. The eastward velocity at 30ЊW at the end of a 30-yr FIG. 11. The baroclinic component of the zonal velocity at 10ЊW 4 4 integration of the ⁄3Њ model experiment ACX. Contour interval: 0.2 of the ⁄3Њ model experiment ACP (taken from the snapshot at the end cm sϪ1. of a 30-yr integration). Contour interval: 0.4 cm sϪ1.

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TABLE 3. The deepest layer interface restored in the GoC, the AC transport at 30ЊW, and the water mass transformation rate in the GoC in the model experiments discussed in the text. The deepest The AC The water mass layer inter- transport transformation The model face restored at 30ЊW rate in the GoC experiment in the GoC (Sv) (Sv)

1 ⁄3Њ model 15 10.50 4.26 ACP 15 7.28 7.13 ACX 13 3.53 2.92 lated horizontally with westward counter¯ow centered at about 36.5ЊN (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 FIG. 12. The baroclinic component of the zonal velocity at 10ЊW 1 as the source and sink in the model is not isolated but of the ⁄3Њ model (taken from the instantaneous ®eld in mid-July of situated against the eastern boundary. The theory pre- the 20th year). Contour interval: 1 cm sϪ1. dicts the upper limit of the transport of the zonal ¯ows and an estimate is given below. Sv), is much lower than the prediction, an indication Following Eq. (7.4.4) of Pedlosky (1996, p. 407), the that the presence of the boundary has a strong impact transport of the eastward ¯ow in the upper ocean (the on the strength of the recirculation. The effect is even AC) may be given as 4 larger in the ⁄3Њ model where the recirculation is absent. yxMW x W f It is likely that the different resolutions mean that details yM TAC ϭ wdxdy0 ϩ wdx0 |,y of the restoring are different (e.g., the number of grid ͵͵ ͵ ␤ S yxSE x E points within the restored area and along the coastline). where w 0 is the vertical velocity (positive downward) Further investigations and experiments are necessary to estimated from the downward mass ¯ux from the model understand the sensitivity of the zonal ¯ows to the de- upper layers to model layer 15 induced by the restoring tails of the boundary forcing, which will be the subject forcing (w 0 is zero outside the restored area); f is the of future research. Coriolis parameter and ␤ is the meridional gradient of f; x and x are the longitudes bounding the restored E W d. The cyclonic cell in the Gulf of Cadiz area; yS is the latitude of the southern edge of the restored area; and yM is the latitude within the restored In the last section, it is shown that the downward area where the eastward ¯ow in the ocean upper layer mass ¯ux induces a cyclonic circulation in the upper 1 reduces to zero or TAC reaches a maximum value, at ocean in the GoC in the ⁄3Њ model. This cyclonic cir- approximately 35.5ЊN in Fig. 12. culation is much enhanced by a strong barotropic com- The ®rst term in the equation is the contribution from ponent as shown in Fig. 13. The vertically integrated the downward mass ¯ux integrated over the area bound- transport in the cell reaches 20 Sv. This barotropic com- ed by xE, xW, yS, and yM, it is estimated to be 2.47 Sv ponent is primarily caused by the use of surface-ref- (if integrated over the entire restored area, it equals to erenced potential density (␴ 0) as explained below. the total water mass transformation of 4.26 Sv). The The use of ␴ 0 means that water masses at depth cannot second term is the contribution from the horizontal re- be properly represented. This is particularly true for the circulation (not the strength of the recirculation) and is Antarctic Bottom Water (AABW) in the deep ocean. the difference between the zonal integral evaluated at Although it is denser than the North Atlantic Deep Water the two latitudes, yS and yM. Note that w 0 is zero at yS, (NADW), which ¯ows above, it has a smaller ␴ 0 value so the second term in this case is the zonal integral at than the NADW due to the compressibility effect. As yM and is estimated to be 40.33 Sv. Thus the eastward ␴ 0 is used as the vertical coordinate in the model, which transport (TAC) predicted by this theory is 42.80 Sv, needs to be monotonic with depth, the AABW and approximately 10 times of the total downward mass ¯ux NADW of the initial ocean state need to be convectively (4.26 Sv) in the GoC. adjusted. As a result, the signature of the AABW is 1 The transport of the AC in the ⁄3Њ model (10.50 Sv), missing in the model, and so is the northward transport although it exceeds the total mass transformation (4.26 associated with it. The transport at the model southern

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1 FIG. 13. The mean barotropic streamfunction (Sv) of the ⁄3Њ model. Contour interval: 5 Sv. boundary due to the restoring condition is shown in Fig. The large southward transport in the bottom layers at 1 14 for the ⁄3Њ model. There is northward transport of the southern boundary has led to the thinning of these light water masses and southward transport of dense layers (in particular, densities 27.88 and 27.92, model water masses, but the northward transport of AABW is layers 15 and 16) during the course of the integration, 4 missing. Such a pattern is also true for all the ⁄3Њ model which in turn, has resulted in the sinking of layer in- experiments presented earlier. terfaces 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 over¯ows 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 expe- rience sinking, which will result in the thickening of the layer just below the deepest restored interface. This is 1 precisely what happens in the ⁄3Њ model in the GoC. The density distribution at 10ЊW is shown in Fig. 15 as an example. There are inversions of ␴ at about 1000 FIG. 14. The transport due to the restoring condition at the southern 0 boundary in classes of model layer densities. A positive value in- m and below 2000 m in the Levitus climatology (Fig. dicates a source (a net northward transport). 15a). These inversions are removed when the dataset is

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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 33.50Њ and 37.50ЊN and above the density contour 27.82 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 4 in the ⁄3Њ 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. 1 16 for the ⁄3Њ 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 over¯ow in the GoC.

←

FIG. 15. The density distribution at 10ЊW: (a) Levitus climatology, 1 (b) the initial state of the ⁄3Њ model, and (c) at the end of a 20-yr 1 integration of the ⁄3Њ model. The density is contoured at model layer densities with two additional contours for 27.69 and 27.90.

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Observations show that the transport of Mediterranean communication). This model uses a restoring condition out¯ow water through the Strait of Gibraltar increases in the GoC to represent the Mediterranean over¯ow pro- signi®cantly as it descends the continental slope by en- cess. Smith et al. (2000) also show an improved GS training the overlying NACW. This entrainment process separation and a realistic AC in a North Atlantic model 1 creates a sink at the eastern boundary for the ocean at ⁄10Њ resolution using the Parallel Ocean Program mod- upper layer in addition to the in¯ow into the Mediter- el developed at Los Alamos National Laboratory. This ranean. Existing theories (Luyten and Stommel 1986; model includes the Western Mediterranean Sea. Anal- Pedlosky 1983, 1996) suggest that a sink of this kind ysis of the results from these models should provide is capable of inducing zonal ¯ows. valuable information on the way in which the GS sep- This mechanism is con®rmed by numerical experi- aration and the Mediterranean over¯ow affect the struc- ments performed with and without the representation of ture of the AC and its associated variability. the Mediterranean over¯ow process. The model does The presence of the AC in response to the restoring not include the Mediterranean over¯ow explicitly, but condition imposed in the GoC suggests that the Medi- restores the model density ®elds in the GoC toward the terranean over¯ow is not only a source of warm and observations. This restoring condition is shown to pro- saline water at depth, but also has a strong dynamic duce a reasonable representation of the water mass trans- impact on the upper-layer circulation in the subtropical formation deduced from observations. eastern North Atlantic. A similar relationship also exists The AC transport is found to be comparable with the between the NAC and the over¯ows across the ridges water mass transformation rate in the GoC in the coarse connecting Greenland, Iceland, and Scotland. BoÈning et resolution model experiments. The eastward ¯owing AC al. (1996) found that in the absence of the over¯ow into the boundary is returned as a westward ¯ow at depth across the Iceland±Scotland Ridge the NAC develops a through the downward mass transfer imposed by the tendency to ¯ow north toward Denmark Strait west of restoring condition in the GoC. This circulation pattern the Mid-Atlantic Ridge to be converted to Denmark is consistent with the eastern boundary ventilation the- Strait over¯ow water. This results in an unrealistic cir- ory of Pedlosky (1983). At the eddy permitting reso- culation pattern in the eastern North Atlantic. The path lution, the AC transport is larger than the water mass of the NAC improves with respect to its eastward pen- transformation rate in the GoC. This enhancement of etration across the Mid-Atlantic Ridge when suf®cient the AC transport comes from the presence of an addi- ¯ow exchange is allowed across the Iceland±Scotland tional horizontal recirculation. Such a recirculation is Ridge (Semtner and Chervin 1992; Roberts and Wood predicted by the theory of dipole circulation associated 1997; Redler and BoÈning 1997). These studies dem- with the isolated sources of Pedlosky (1996) although onstrate that over¯ows as a whole in the North Atlantic the strength of the circulation is much weaker in the make a strong impact on the large-scale circulation in model due to the presence of the boundary. This analysis the upper ocean and emphasize that they need to be suggests that the zonal ¯ows are sensitive to the details properly represented in ocean general circulation mod- of the forcing at the boundary. els. Further, over¯ows at high latitudes of the North Other factors, though not investigated, may play a Atlantic form an important part of the thermohaline cir- part in shaping the model AC. For example, the position culation of the World Ocean (Broecker 1991), which of the GS separation may have in¯uenced the separation may have signi®cant effects on climate variability (Del- of the GS into the NAC and the AC in the southeast of worth et al. 1993). the Grand Banks. The model GS separates from the coast too far north due to the insuf®cient resolution to Acknowledgments. I am very grateful to Beverly de account for the complex western boundary processes Cuevas and David Webb for their helpful comments on (OÈ zgoÈkmen et al. 1997). the draft manuscript. My sincere thanks go to the two It seems reasonable to speculate that the combined anonymous reviewers who provided the most thorough 1 effect of western boundary processes (which govern the and critical reviews of the manuscript. The ⁄3Њ version GS separation, its path downstream and its separation of the model was developed and integrated under the into the NAC and the AC) and the Mediterranean over- DYNAMO project funded by the EU MAST Contract ¯ow process (which induces and maintains the AC and MAS2-CT93-0060. The contributions from Sally Bar- possibly encourages the separation of the GS into the nard and Adrian New to the project are much appre- NAC and the AC) largely determines the observed AC ciated. pathway. Thus it is essential to have these two processes properly represented before a realistic simulation of the REFERENCES AC can be achieved in general circulation models. Recent developments in high resolution modeling are Ambar, I., and M. R. Howe, 1979: Observations of the Mediterranean promising in these respects. For instance, the high res- out¯ow. II. The deep circulation in the vicinity of the Gulf of Cadiz. Deep-Sea Res., 26A, 555±568. olution simulation of the North Atlantic performed with Baringer, M. O., and J. F. Price, 1997: Mixing and spreading of the 1 MICOM at ⁄12Њ resolution shows an improved GS sep- Mediterranean out¯ow. J. Phys. Oceanogr., 27, 1654±1677. aration and a realistic AC (E. Chassignet 1998, personal Barnier, B., L. Siefridt, and P. Marchesiello, 1995: Thermal forcing

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