The Red Sea outflow regulated by the Indian monsoon

Continental Shelf Research 26 (2006) 1448 1468 www.elsevier.com/locate/csr The Red Sea outflow regulated by the Indian monsoon Hidenori Aiki a,, Keiko Takahashi b, Toshio Yamagata a,c a Frontier Research Center for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokohama 236-0001, Japan b Earth Simulator Center, Japan Agency for Marine-Earth Science and Technology, Yokohama 236-0001, Japan c Department of Earth and Planetary Science, University of Tokyo, Tokyo 113-0033, Japan Available online 30 June 2006 Abstract To investigate why the Red Sea water overflows less in summer and more in winter, we have developed a locally highresolution global OGCM with transposed poles in the Arabian peninsula and India. Based on a series of sensitivity experiments with different sets of idealized atmospheric forcing, the present study shows that the summer cessation of the strait outflow is remotely induced by the monsoonal wind over the Indian Ocean, in particular that over the western Arabian Sea. During the southwest monsoon (May September), thermocline in the Gulf of Aden shoals as a result of coastal Ekman upwelling induced by the predominantly northeastward wind in the Gulf of Aden and the Arabian Sea. Because this shoaling is maximum during the southwest summer monsoon, the Red Sea water is blocked at the Bab el Mandeb Strait by upwelling of the intermediate water of the Gulf of Aden in late summer. The simulation also shows the three-dimensional evolution of the Red Sea water tongue at the mid-depths in the Gulf of Aden. While the tongue meanders, the discharged Red Sea outflow water (RSOW) (incoming Indian Ocean intermediate water (IOIW)) is always characterized by anticyclonic (cyclonic) vorticity, as suggested from the potential vorticity difference. r 2006 Elsevier Ltd. All rights reserved. Keywords: Stretched-coordinate OGCM; Monsoon winds; Coastal upwelling; Subsurface tongue 1. Introduction The Red Sea is a semi-enclosed mediterranean sea surrounded by the African and Eurasian continents, which is linked to the Indian Ocean by a very shallow sill. The excess evaporation over the Red Sea produces extremely salty and dense water which intermittently spills from the sill and cascades down to the intermediate depths of the Indian Ocean. Passing through the Gulf of Aden (Fig. 1), the Red Sea outflow water (RSOW) becomes a part of the Corresponding author. E-mail addresses: aiki@jamstec.go.jp (H. Aiki), yamagata@eps.s.u-tokyo.ac.jp (T. Yamagata). intermediate circulation in the Indian Ocean, which has been observed as a mid-depth salinity maximum in the Arabian Sea and even in the southern hemisphere (Wyrtki, 1971; Quadfasel and Schott, 1982; Gordon et al., 1987; Beal and Bryden, 1997; Mecking and Warner, 1999). Compared with the other marginal overflows in the world oceans (e.g. Mediterranean water outflow in the North Atlantic), the Red Sea outflow system has several unique features: (i) the outflow from the Red Sea ceases in late summer, (ii) the strait connecting the Red Sea to the Gulf of Aden is very steep and narrow, and (iii) the Gulf of Aden is dominated by a chain of mesoscale eddies interacting with the tongue of RSOW underneath. 0278-4343/$ - see front matter r 2006 Elsevier Ltd. All rights reserved. doi:10.1016/j.csr.2006.02.017

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1449 Fig. 1. (a) Bathymetry of the region from the Red Sea to the Indian Ocean. Contours show bottom depths of 100, 500, and 1000 m. (b) Curvilinear orthogonal coordinates used in the present numerical model. (c) The global map of the transformed coordinates in Mercator projection, with the upper (lower) end corresponding to the transposed pole in the Arabian peninsula (India). (d) The close-up topography of the numerical model around the Bab el Mandeb Strait. Contours show bottom depths of 150, 200, 300, 600, and 800 m. The Bab el Mandeb Strait is located at the exit of the Red Sea and consists of several complex topographic features. From the Red Sea to the Indian Ocean along the 200-km-long Strait, there are the Hanish Sill, the Perim Narrows, and two bottom outlets connected to the Tadjura Rift (Fig. 1). The Hanish Sill is only 160 m deep and 5 km wide at the bottom, so that it is crucial in determining the nature of RSOW. At the tip of the southwestern corner of the Arabian peninsula is the Perim Narrows of about 20 km wide, in which the strait deepens to 300 m. As the strait further slopes

1450 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 down to 600 800 m in the southern subsection of the strait, it separates into two bottom branches called the Northern and Southern Channels. These tiny trenches (not visible in Fig. 1a) are curved on the flank of the Tadjura Rift of 1600 m deep. The main outlet of RSOW is the 120-km-long Northern Channel (Peters et al., 2005; Peters and Johns, 2005), which is a narrow bottomed valley of only 5 km wide. As the descending RSOW is sheltered by the confined topography of the Northern Channel, the bottom salty water does not experience significant mixing with the ambient water until it is ejected into the mid-depths of the Tadjura Rift. It is very challenging for a numerical model to resolve the above described topography of the Bab el Mandeb Strait. O zgo kmen et al. (2003) used a two-dimensional nonhydrostatic model to partly investigate the sinking of RSOW through the Northern Channel. The present study uses a threedimensional stretched-coordinate model that incorporates the narrow and steep topography of the Bab el Mandeb. At the western end of the Gulf of Aden, the discharged RSOW reaches a neutral-buoyancy depth of 400 800 m (isopycnals of 27.0 27:4s y ) and leaves the coast to propagate eastward (cf. Bower et al., 2000). The excursion of RSOW is traceable as a mid-depth salinity maximum in the Gulf of Aden and further downstream in the Indian Ocean (Warren et al., 1966; Quadfasel and Schott, 1982; Meschanov and Shapiro, 1998; Mecking and Warner, 1999; Han and McCreary, 2001). Recent observations by Bower et al. (2002) have revealed the meandering pathways of RSOW produced by interaction with a chain of deep-reaching mesoscale eddies in the Gulf of Aden. Due to the lateral stirring by the mesoscale eddies, the RSOW undergoes considerable (isopycnal) mixing in the Gulf of Aden. The well-developed core of the RSOW enters the Arabian Sea and sheds lenses of salty water, which are often called Reddies (Red Sea water eddies). In summarizing the distribution of the observed salinity anomalies in the Arabian Sea, Shapiro and Meschanov (1991) suggest that the most probable mechanism of Reddy generation is instability of the RSOW main tongue. The numerical simulations in the present study complement sparse measurements in the northwestern Indian Ocean, and allow us to understand the three-dimensional structure of the RSOW tongue (Section 3). At the Bab el Mandeb Strait, the salty Red Sea water is discharged mainly in winter and it ceases in Table 1 The annual mean values, together with the winter (November May) values over the summer (June October) values in the parentheses, of observed transport Q obs (Sv) salinity S obs (psu), temperature T obs (1C), and potential density r obs ðs y Þ through the Bab el Mandeb Strait Q obs (Sv) S obs (psu) T obs (1C) r obs (s y ) Surface layer 0.31 0:54 0:01 37.0 36:6 37:3 28.3 26:7 30:5 23.8 24:0 23:7 Middle layer 0.07 0:18 36.7 36:7 24.4 24:4 24.8 24:8 Bottom layer 0.36 0:52 0:14 39.7 39:6 39:8 22.8 23:3 22:1 27.6 27:3 27:8 These are calculated from the monthly mean values given in Table 7 of Sofianos et al. (2002). A positive (negative) Q denotes an outflow (inflow). The surface, middle, and bottom layers are defined by the zero crossings of the observed velocity profile.the middle layer does not exist in the winter season, which is denoted by the. summer. The winter exchange (about 0.5 Sv in Table 1, 1Sv¼ 10 6 m 3 /s) occurs in the surface and bottom layers, whereas the summer exchange (about 0.16 Sv in Table 1) occurs in the middle and bottom layers (Thompson, 1939; Murray and Johns, 1997). During the summer monsoon (May September), some water mass from the intermediate layer of the Gulf of Aden (i.e. the Indian Ocean) intrudes into the Red Sea (Jones and Browning, 1971; Patzert, 1974; Be thoux, 1987; Maillard and Soliman, 1986; Murray and Johns, 1997; Saafani and Shenoi, 2004). It has been generally believed (Thompson, 1939) that the above-mentioned seasonal cycle of the water exchange is driven primarily by the local wind stress which is constrained by the orography of the Bab el Mandeb. The predominant winds over the strait blow towards the Red Sea during the winter monsoon (October April) to enhance the surface layer inflow (and bottom outflow), whereas the wind blows towards the Indian Ocean during the summer monsoon (May September) to stop the surface layer inflow (and bottom outflow). In contrast, the water exchange through the Bab el Mandeb appears to be hydraulically controlled (Pratt et al., 1999; Smeed, 2000, 2004; Siddall et al., 2002). This highlights the stratification in the upstream and downstream of the strait: upwelling in the Gulf of Aden during summer may play an important role for the seasonal outflow and inflow (Jones and Browning, 1971; Patzert, 1974; Smeed, 2000). The relative importance of these two wind-forced effects, one direct and one indirect, is not clear. To examine this issue, we will present in Section 4 a series of sensitivity

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1451 experiments with different sets of idealized atmospheric forcing. This paper is organized as follows. Section 2 introduces a locally high-resolution OGCM for the Red Sea and Gulf of Aden region, which is integrated for 15 years in our reference experiment. We investigate in Section 3 the three-dimensional structure and time evolution of the RSOW tongue in the Gulf of Aden. Section 4 presents a series of sensitivity experiments demonstrating that the seasonal overflow through the Bab el Mandeb is remotely constrained by the monsoonal currents in the Indian Ocean. This paper ends with a brief summary in Section 5. 2. Numerical model We have developed a stretched-coordinate OGCM focusing on the Red Sea and the Gulf of Aden region (Fig. 1). The present configuration is designed to incorporate the interactions between the small-scale processes around the Bab el Mandeb Strait and the larger-scale currents in the Indian Ocean. The computational domain covers the entire globe, and hence there are no open boundaries. 2.1. Stretched coordinates The present model uses the curvilinear orthogonal coordinates given by the pole-transposition method of Bentsen et al. (1999). The region covering the Red Sea and the Gulf of Aden is focused on by locating one pole at the southwestern corner of the Arabian peninsula (10 km inside the coastline, closest to the Perim Narrows) and another pole in India (Fig. 1b). As shown in Fig. 1c (with the Mercator projection), the transformed coordinates deform the world oceans to enlarge the Arabian marginal seas; about half of the computational domain is occupied by the analysis region of the present study. This transformed map is discretized on a 240 256 grid in the present model, giving a smallest horizontal grid spacing of about 2 km around the narrowest region of the Bab el Mandeb. The Perim Narrows (about 20 km wide at the sea surface) are well resolved, but the resolution is insufficient to capture realistically the steep and narrow topography at the bottom of the strait, in particular the Northern Channel that is 5 km wide at the outlet to the Gulf of Aden (see Section 1). The grid spacing increases from 4 to 20 km in the Gulf of Aden (Fig. 7a), and is capable of capturing the mesoscale eddies (scales 100 km) reported by Bower et al. (2002). As the resolution around the head of the Somali Peninsula is about 20 km, it is sufficient for simulating the Great Whirl and the Socotra Gyre, along with the seasonal evolution of the Somali Current along the African coast (Swallow and Bruce, 1966; Warren et al., 1966; Schott, 1983; Schott et al., 1997; Esenkov and Olson, 2002; Wirth et al., 2002). In the interior of the Indian Ocean, the large-scale currents driven by monsoonal winds can be resolved with the average grid spacing of about 80 km (cf. Schott and McCreary, 2001; Shankar et al., 2002). The Pacific, Atlantic and Southern Oceans are included at low resolutions of 200 400 km and represent almost passive absorbing boundaries. The CFL condition is very severe at the finest mesh around the Bab el Mandeb, giving a time step of 1.5 min in the numerical integration. Therefore the computational cost of integrating the present model with 61 440 ð¼ 240 256Þ horizontal grid points for 15 years is comparable to that of integrating a 11 global OGCM with 64 800 ð¼ 360 180Þ grid points for 150 years with a time step of 15 min. The vertical grid of the present model is given by a hybrid terrain-following and z-level coordinate, called the stepped sigma coordinate (Hanney, 1991; Beckmann and Haidvogel, 1993; Ezer and Mellor, 2004). This coordinate has the advantage of packing the numerical mesh near the sloped bottom without causing a significant error in the pressure gradient. The grid spacing is the smallest (8 m) at the sea surface, which monotonically increases to the bottom with 30 elements. As a result, the Hanish Sill (160 m deep) in the Bab el Mandeb is grided uniformly with the minimum vertical spacing (of 8 m), whereas at a mid-depth of 800 m in the Gulf of Aden (where parcels of RSOW drift) the vertical resolution is about 80 m; the vertical grid distribution is shown in Figs. 6, 9b, 10b. The model topography of the sea floor is constructed from the 1-min global bathymetric dataset of the General Bathymetric Chart of the Oceans. Over the regions of the Southern and Northern Channel, we needed to modify the topographic data to represent the steep outlet to the Gulf of Aden as shown in Fig. 1d. The model Northern Channel (between point A and BinFig. 1d) lowers the bottom depth from 250 to 900 m. In contrast to the real bathymetry (e.g. Figure 1 of Peters et al., 2005 and Figure 1c of Bower et al., 2005), the route of the model Northern

1452 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 Channel is almost straight. The model Channel has somewhat taller side walls than that in the real topography, as we discuss later in Section 3.2. The Southern Channel is not included in the present model, so that the preset model is unable to simulate the multi-layer structure of RSOW above the Tadjura Rift (cf. Bower et al., 2005). The present study is more concerned with the main body of RSOW at intermediate depths 500 900 m in the Gulf of Aden, and its interaction with currents in the northwestern Indian Ocean. 2.2. Formulation and forcing The strong tidal effects near the Perim Narrows and the Hanish Sill may change the characteristics of the strait water (Jarosz et al., 2005a,b). Thus the present model includes the tide-generating potential of its major four constituents (M 2 ; S 2 ; K 1, and O 1 ) and the nonhydrostatic pressure formulation of Marshall et al. (1997). A prototype of the present model was used in Aiki and Yamagata (2004) to simulate the shedding of Meddy-like lenses from the Mediterranean Sea. The bottom friction is calculated from a standard quadratic equation. The horizontal eddy viscosity is given by Smagorinsky (1963) with a nondimensional coefficient of 0.1 (after squaring). The vertical eddy viscosity is set 100 cm 2 =s ð1cm 2 =sþ when the Richardson number is smaller (larger) than 0.5 to parameterize either Kelvin Helmholtz or Holmboe s instability. The Prandtl number is unity near the surface (shallower than 100 m) as the upper and surface mixed layers are somewhat diabatic, whereas the Prandtl number is set to 10 in the subsurface layer below 100 m depth to avoid excessive diffusion of RSOW. O zgo kmen et al. (2003) adopted a large Prandtl number of 20 in their nonhydrostatic numerical experiments. The present model adopts no special measure for the surface mixed layer and the bottom boundary layer. The flux-corrected transport scheme (Zalesak, 1979) is used for the advection of tracers. The temperature and salinity fields of the World Ocean Atlas 1998 are used for the initial condition with no motion: the velocity field is later determined by geostrophic adjustment during the initial steps of time integration. The model is integrated for 15 years with the atmospheric forcing from monthly ECMWF (European Centre for Medium Range Weather Forecast) reanalysis wind stress data and COADS (Coastal Observational And Data Ship) water and heat fluxes dataset. There is no restoration to the temperature and salinity climatologies at the sea surface. As in Sofianos and Johns s (2002, 2003) numerical study on the Red Sea circulation, we correct the COADS flux data over the Red Sea by adding 0.5 m/year for evaporation precipitation and 60 W/m 2 for surface heat flux, respectively. These values are based on Sofianos et al. s (2002) estimates of the annual mean freshwater and heat fluxes over the Red Sea: these are 2 m/year and 11 W/m 2 whereas the COADS values are 1.5 m/year and 40 W/m 2, respectively. The Knudsen formulae used in Sofianos et al. (2002) dictate the freshwater and heat budgets inside the Red Sea: r 1 Q 1 S 1 þ r 2 Q 2 S 2 þ r 3 Q 3 S 3 ¼ 0, (1) ðe PÞA ¼ Q 1 þ Q 2 þ Q 3, (2) F T ¼ 1 A r 0c p ðq 1 T 1 þ Q 2 T 2 þ Q 3 T 3 AðE PÞT S Þ, (3) where r i, Q i S i, and T i are the density, layer transport, salinity, and temperature for the surface water ði ¼ 1Þ, the intermediate water ði ¼ 2Þ, and the bottom water ði ¼ 3Þ through the Bab el Mandeb; E P is the fresh water flux over the Red Sea; A is the surface area of the Red Sea; F T is the heat loss over the Red Sea; r 0 is the reference density ð1026 kg/m 3 Þ; c p is the heat capacity of water ð3986 J C=kgÞ; T S is the Red Sea surface temperature. The overbar denotes the annual mean. To calculate E P or F T, Sofianos et al. (2002) used the values of Q i, T i, and S i given by Murray and Johns (1997) except for the surface layer transport Q 1 (due to the large error in the observational estimates). Note that the above correction is only for the annual mean budget. More detailed correction including seasonal and spatial variations is necessary but it is left for a future study (cf. Cromwell and Smeed, 1998). 3. Results of the reference experiment Throughout the 15 years of simulation for the reference experiment, we released a passive tracer representing the concentration of RSOW from the Red Sea. This tracer is in a statistically equilibrium state during the last 5 years of the reference experiment (Fig. 2), which also suggests a mean resident time of about 2.3 years for RSOW in the Gulf of Aden. Model results from the final year are

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1453 total amount (a) total amount (b) RSOW inside Gulf RSOW over the globe year Fig. 2. Total amount of the passive tracer (a) over the Gulf of Aden and (b) over the entire region outside the Red Sea, during the last 5 years of the reference experiment. Both ordinates are nondimensionalized by 6:3072 10 15 m 3 ð¼ 0:22 Sv 86400 s 365 dayþ: an annual discharge amount of the RSOW tracer, calculated from the bottom figure (i.e. the slope of the line in the bottom figure is made unity). shown in the followings to discuss the development of the subsurface and surface flows in the northwestern Indian Ocean. Sections 3.1 and 3.2 will serve to validate the model results with observations. Sections 3.3 and 3.4 focus on the threedimensional distribution of the RSOW tracer in the Gulf of Aden. 3.1. Currents at the sea surface Upper layer currents around the Gulf of Aden are displayed with the vertical component of relative vorticity at the sea surface in Fig. 3. The simulated life-cycle of the Somali Current system is consistent with previous reports (Swallow and Bruce, 1966; Warren et al., 1966; Schott et al., 1997; Schott and Fischer, 2000; Esenkov and Olson, 2002; Wirth et al., 2002). The southwest monsoon (May September) in the Indian Ocean drives the Somali Current to flow northward along the African coast (Figs. 3a,b), which develops into the Great Whirl and Socotra Gyre in late summer (Fig. 3c). The Rossby number of these intense currents and eddies is large (above 0.5), attributable in part to the small Coriolis parameter at low latitudes. The onset of the northeast monsoon (October April) results in the reversal of the Somali Current (Fig. 3d). It is interesting that the Gulf of Aden is filled with an array of mesoscale eddies throughout the year; these are the Gulf of Aden eddies reported in Bower et al. (2002). These Gulf of Aden eddies are highly baroclinic with respective to the main thermocline at a depth of 300 m and the lower part extends to the bottom, as we explain in Section 3.4 (Fig. 10). Over the Bab el Mandeb Strait in Fig. 3, an energetic inflow to the Red Sea is found in all months except September when the surface shows an outflow to the Gulf of Aden. The simulated reversal of the surface current is consistent with the observation of the surface layer outflow during July September (Sofianos et al., 2002). The seasonal contrast of the water exchange through the strait is better exhibited in the sea surface salinity. In March (Fig. 4a) there is an inflow of fresher water from the Indian Ocean that extends to 171N inside the Red Sea. In August (Fig. 4b) there is a counter current of salty Red Sea water along the African coast, whose outgoing edge forms an anticyclonic eddy in the western gulf. Although the local wind stress over the Bab el Mandeb Strait is underestimated in the present study, the surface current through the Bab el Mandeb Strait is shown to reverse seasonally. The bottom outflow at the Bab el Mandeb will be detailed in the next subsection. 3.2. The Bab el Mandeb Strait The reference experiment is successful in reproducing the seasonal transition of the water exchange through the Bab el Mandeb Strait. The along-strait velocity at the Perim Narrows (Fig. 5a) shows a two-layer exchange from October to June and a three-layer exchange during late summer. The bottom outflow has a maximum speed of 120 cm/s during March April, which is comparable to the observed speed in Figure 3 of Murray and Johns (1997) and Figure 3 of Sofianos et al. (2002). In order to estimate the volume transport of the bottom outflow, one needs to define the upper interface of the bottom water. We have first calculated a cross-strait integral of the along-strait velocity at each depth in the cross-section. Then, as in Sofianos et al. (2002), zero-crossing in the vertical profile of the each-depth transport ðm 2 =sþ is used to estimate the overflow rate. The reference experiment (see solid line in Fig. 11) gives a winter maximum of 0.51 Sv and a summer minimum of 0.11 Sv with an

1454 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 Fig. 3. Seasonal evolution of relative vorticity (RV; the vertical component) at the sea surface during the 15th year of the reference experiment. Positive (negative) values indicate cyclonic (anticyclonic) rotation, with units of the Coriolis parameter. annual average of 0.36 Sv. Thus, even using the large-scale wind forcing as in the present study, we are able to reproduce a seasonal difference of 0.4 Sv in the overflow rate of RSOW. The simulated seasonal difference of 0.4 Sv is greater than both theoretical and model estimates in Siddall et al. (2002) and Sofianos and Johns (2002). The overflow rate of the present study compares qualitatively with the observations (Table 1), although the direct observations of Murray and Johns (1997) showed a short-term peak of 0.7 Sv in February and a complete cessation of bottom overflow in August. The discrepancies are on short time scales ð1 monthþ and are possibly due to the limited wind and buoyancy forcing applied in the present model (Section 2.2). The lack of high-frequency (and also locally intensified) wind in the present simulation may be one of the reasons why the middle layer velocity (during July September) does not reach the bottom in Fig. 5a. Temperature profile at the Perim Narrows highlights intrusion of the Gulf of Aden intermediate water in late summer. In Fig. 5b, a minimum temperature of 18 1C is found at 160 m depth during

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1455 Fig. 4. Sea surface salinity (SSS, [psu]) around the Bab el Mandeb Strait in (a) March and (b) September, simulated in the 15th year of the reference experiment. August September. Indeed such cold intrusion has been observed by Saafani and Shenoi (2004). The cold signal of the Gulf of Aden intermediate water continues northward across the Hanish Sill where the core is warmed to about 22 1C (not shown). Also comparison between the temperature in Fig. 5b and that in Table 1 is meaningful. For example in August in Fig. 5b, if one focuses on values of temperature at the depths of the velocity maximum, one finds 31 1C at the sea surface, 23 24 1C at 60 m depth, and 22 1C at 190 m depth. These temperature values are within an accuracy of 1 C if compared with the values of the surface, middle, and bottom layer in Table 1 for summer season; see Table 2 for a summary of the simulated temperature, salinity, and density profiles at the Perim Narrows. In Figs. 5b d, all of temperature, salinity and density profiles are vertically homogenized in the vicinity of the upper interface of the bottom water, which results from enhanced diabatic mixing inside the Strait. While the diabatic mixing allows detrainment of RSOW to overlying fluids, bottom salinity remains greater than 39.5 psu throughout the year (Fig. 5c). The intrusion of the Gulf of Aden water is also exhibited by uplifted isopycnals during July September in Fig. 5d. Again, if salinity and density values in Fig. 5 are picked up at the depth of velocity maximum at any time in (see Table 2), the simulation result provides a good comparison to the observed values in Table 1. Once being ejected from the Perim Narrows, most parcels of RSOW enter the model Northern Channel and cascade downslope. As has been observed by Peters et al. (2005), the bottom salinity is greater than 39.5 psu down to 500 700 m depth in March when the overflow of RSOW is strongest (Fig. 6a). However, the overlying fluids are unrealistically simulated in salinity range between 39.0 and 39.5 psu, distributing up to 250 m depth and over the along-channel distances between 40 and 100 km. This thick layer of the detrained RSOW is shaped by the sidewalls of the model Northern Channel which is somewhat taller than the real one: the wall height is 400 m at an along-channel distance of 100 km in Fig. 6, whereas the height in the real topography is less than 300 m at Station 37 in Figure 1a of Peters et al. (2005). In August (Fig. 6b), the bottom core of RSOW retreats from the Channel leaving behind water parcels of decreased salinity. In summary, the present simulation is able to reproduce bottom salinity and seasonal changes of RSOW in the Northern Channel. The model bias at 300 500 m depth will not be serious for the analysis of the RSOW tongue distributing below 500 m depth in the offshore region. Hopefully the excessive dilution at the upper depths compensates the lack of the Southern Channel in the present model. 3.3. Spreading of RSOW The Gulf of Aden is filled with random energetic eddies (Fig. 3) and hence the distribution of RSOW incorporates year-to-year variability. Results explained in the following are an example for one

1456 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 Perim velocity Table 2 Same as Table 1 except for the result of the reference experiment depth [m] S model (psu) T model (1C) r model ðs y Þ Surface layer 36.6 36:4 37:1 28.1 27:2 30:9 23.5 23:7 23:0 Middle layer 36.8 36:8 25.5 25:5 24.5 24:5 Bottom layer 39.0 39:0 39:0 23.2 23:5 22:1 26.9 26:8 27:2 (a) Perim temperature Salinity, temperature, and density values in Fig. 5 are picked up by using the transport-weighted-mean as in Sofianos et al. (2002). The model s winter (summer) mean values are from October to June (July September). depth [m] (b) depth [m] (c) depth [m] (d) Perim salinity Perim density Fig. 5. Vertical profiles at the Perim Narrows for (a) along-strait velocity (cm/s), (b) temperature (1C), (c) salinity (psu), and potential density ðkg=m 3 Þ, averaged over the last 3 years of the reference experiment. These four quantities are taken at a deepest point in the cross-section (point A in Fig. 1d), corresponding to the measurements of Murray and Johns (1997) and Sofianos et al. (2002). Positive (negative) values in Fig. 5a indicate an outflow (inflow). Shade in each figure shows depths where the each-depth transport ðm 2 =sþ is negative (directed to the Red Sea). particular year. Fig. 7 shows the time evolution of the subsurface distribution of RSOW: the vertical integral of the RSOW concentration (the passive tracer) is interpreted as an effective thickness of RSOW. After being discharged from the Bab el Mandeb, RSOW initially piles up in the Tadjura Rift (the western edge of the Gulf) with a thickness of 700 m. In January (Fig. 7a), RSOW forms two large blobs (about 120 m thick) in the western Gulf of Aden. The main blob centered at (461E, 121N) spans the meridional extent of the gulf, and is associated with anticyclonic rotation exhibited in Fig. 7g. This large anticyclone is coupled with a small cyclone to the west at (44.71E, 121N), which consists of a patch of low concentration RSOW: the mixed product of RSOW and Indian Ocean Intermediate Water (IOIW). The second blob of RSOW is located along the African coast at (44.51E, 111N) in Fig. 7a, which is part of the southward-flowing branch of newly ejected RSOW wrapping around the cyclonic vortex mentioned above. This boundary stream of RSOW appears to leave the African coast at 451E to continue to the above-mentioned anticyclonic blob of RSOW. Meanwhile, the eastern half of the Gulf of Aden contains a thin and large (about 400 km long and 150 km wide) tongue of RSOW that is associated with the negative (anticyclonic) relative vorticity (Fig. 7g). This indicates that the RSOW tongue is here on the south of an outgoing current to the Indian Ocean as is a front comprising RSOW and IOIW. In March (Fig. 7b), the RSOW tongue (100 m thick) in the eastern gulf slightly rotates its axis to the south while exhibiting a clear pair of relative vorticity (Fig. 7h), so that the intermediate layer current is now directed to Socotra Island. In the western gulf, the aforementioned main anticyclonic blob of RSOW is pushed northward to the Arabian coast. At the western edge of the Gulf, the newly

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1457 A B depth [m] depth [m] March (a) August (b) distance [km] Fig. 6. Vertical distribution of salinity (psu) along the model Northern Channel (between point A and B in Fig. 1d) in (a) March and (b) August, from the 15th year of the reference experiment. Contours show the vertical grid distribution with every five elements. ejected stream of RSOW from the Tadjura Rift remains following the African coast. In May (Fig. 7c), the body of RSOW exhibits a gulf-long stream, which meanders but extends over the whole range of the main axis of the Gulf of Aden. This meandering suggests frontal instability at the subsurface layers (Fig. 7i), which may lead to the isopycnal mixing. The intense anticyclonic rotation found at (461E, 12.51N) is associated with the thick RSOW blob evident in Fig. 7c at the same position. Interestingly in July (Fig. 7d), as a result of the breaking of the RSOW stream, the eastern gulf contains a well-mixed product of RSOW and IOIW of about 100 m thick. On the other hand, the western gulf contains two thick blobs of RSOW of about 200 m thick, with the offshore one at (46.51E, 121N) consisting of the anticyclonic part of a dipolar vortex in Fig. 7j. The end of the summer monsoon brings a drastic change in the Gulf of Aden as most of RSOW parcels shift southward to the African side of the Gulf of Aden (Fig. 7e). Interestingly, a large blob of IOIW is seen on the Arabian side at (481E, 131N) in Fig. 7e, indicating the intrusion of IOIW from the east. Indeed, the time sequence during July November (Figs. 7c e) shows several patches of IOIW translating westward along the Arabian coast. This westward progression replaces the western gulf with IOIW (Fig. 7f), which in turn pushes the main body of the RSOW tongue eastward, towards the Indian Ocean. These movements of intermediate water is part of a large cyclonic circulation at the mid-depth (Fig. 7j). Results from the other years show that the above-mentioned scenario of the replacement of IOIW and RSOW does not apply to all years. Occasionally no major patch of IOIW intrudes into the Gulf of Aden in a year: in some years a small patch of IOIW intrude from the Somalian Coast. Regardless of these year-to-year differences, the total amount of RSOW is in a statistically equilibrium state (Fig. 2a). Detailed investigation is devoted to a future study. The above results have demonstrated that, while the tongue meanders, the discharged RSOW (the incoming IOIW) is always characterized by anticyclonic (cyclonic) vorticity. This may result from the

1458 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 Fig. 7. Time series of the reference experiment during the 15th year. (a) (f) Vertical integral (below 26:5s y ) of the passive tracer released from the Red Sea (which is initially unity), displaying an effective thickness (m) of the Red Sea outflow water (RSOW). (g) (l) Relative vorticity (RV; the vertical component) averaged between isopycnals of 27.0 27:5s y, around a depth of 800 m. Positive (negative) values indicate cyclonic (anticyclonic) rotation, with units of the Coriolis parameter. The contours in panel (a) show the horizontal resolution (km) of the present model.

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1459 difference of potential vorticity, as has been suggested for Meddies in the North Atlantic (Morel and McWilliams, 1997; Aiki and Yamagata, 2004). 3.4. Vertical profiles of RSOW To investigate how RSOW leaves the coast and penetrates into the offshore stratification, we have prepared Fig. 8 which shows a vertical section along the main axis of the Gulf of Aden. The body of the RSOW tongue is distributed at depths between 600 and 1000 m, which is similar to the observation of Bower et al. (2002). The RSOW parcels mainly drift along the 27.3 27:4s y isopycnal in Fig. 8, which is somewhat deeper than the observed isopycnal (27.15 27:35s y ) of the salinity maximum (Meschanov and Shapiro, 1998). The density and depth of the simulated RSOW tongue are highly sensitive to both the model topography of the Bab el Mandeb and the parameterization of diapycnal mixing near the bottom (Ezer, 2005; Legg et al., 2006). Fig. 8 shows two blobs of RSOW in the downstream at 491E and 511E. These are associated with the meandering of the stream (Fig. 7c). The RSOW is subject to enhanced isopycnal mixing by the frontal instability in the Gulf of Aden (where the concentration reduced to 0.2), whereas further upstream inside the strait the RSOW is sheltered by the confined topography and thus experiences less mixing (the concentration of RSOW was 0.6 after being discharged from the strait). North south sections along 44.41E (Fig. 9) and 501E (Fig. 10) are shown for salinity and (the vertical component of) relative vorticity. In the western side of the gulf (Fig. 9a), the discharged high salinity water (over 37 psu) is well above the bottom and is here neutrally buoyant (cf. Bower et al., 2005). The relative vorticity in Fig. 9b exhibits a bottom reaching cyclone associated with the depression of RSOW in Fig. 9a near the northern boundary (around 12.11N). The thermocline is located at 300 m depth above which there is a pair of positive and negative relative vorticity. Fig. 10a, which is another vertical section crossing the tip of the RSOW tongue, displays a small core (about 50 km wide) of salty water drifting at 800 m depth which is surrounded by much diluted RSOW parcels distributed in parallel to the isopycnal surfaces. This confirms the enhanced isopycnal mixing due to the swinging of RSOW tongue. The relative vorticity distribution in Fig. 10b is quite interesting because it is dominated below the thermocline at 300 m depth by a tall vorticity pair whose zero-crossing (at 13.21N) is exactly aligned at the aforementioned RSOW core drifting at the middepth. This tall vorticity pair is associated with the aforementioned eastward (outgoing) current to the Indian Ocean (Figs. 7g i). 4. Sensitivity experiments on the seasonal overflow The following subsections present the results of a series of sensitivity experiments on the last 5 years of Fig. 8. Vertical distribution of the concentration of the Red Sea water (shaded) and the potential density (dotted line) ðkg=m 3 Þ, along the oblique line in Figs. 7c and i. The dotted line shows isopycnal contours.

1460 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 Fig. 9. Vertical distribution of (a) salinity (psu) and (b) relative vorticity (normalized by the Coriolis parameter) along 44.41E of Figs. 7c and i. The dotted line in panel (a) shows isopycnal contours. The solid line in panel (b) shows the vertical grid distribution with every five elements. Fig. 10. Same as Fig. 9 except for the section along 501E. the reference experiment, which we performed to elucidate how and why the Red Sea water overflows seasonally through the Bab el Mandeb. 4.1. Wind or buoyancy forcing? In order to verify whether the seasonal variation in the water exchange through the strait is due to wind stress or atmospheric buoyancy (water and heat) fluxes, we have performed two sensitivity experiments using different forcing. The first run ( wind-run in Fig. 11) is driven by the monthly mean wind stress (as in the reference experiment) but with steady (i.e. annual mean) climatology heat and water fluxes. The result (Fig. 11a) is almost identical to the reference experiment (abbreviated ref-run ), showing the seasonal variation in the overflow rate with a winter maximum of 0.46 Sv and a summer minimum of 0.07 Sv. Another run ( buoyrun in Fig. 11) has used the monthly mean heat and water fluxes (as in the reference experiment) but with a steady (i.e. annual mean) wind stress. This results in little seasonal variability (Fig. 11a): the overflow rate is about 0:38 0:5 Sv throughout the year. From these two results it can be concluded that the summer stopping and the winter culmination of RSOW transport is primarily due to the seasonal evolution of wind stress rather than to atmospheric buoyancy forcing. The dominant role of wind stress is consistent with the previous studies that have attributed the seasonal water exchange to wind stress applied above the Bab el Mandeb (cf. Thompson, 1939). 4.2. Where is the wind stress important? We have next conducted three sensitivity experiments to identify which region of the wind stress is essential to the seasonal variability in the strait water exchange. In the first experiment ( red-run ), the monthly wind is applied only over the Red Sea (north of 141N), with other regions receiving the

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1461 overflow rate [Sv] overflow rate [Sv] overflow rate [Sv] (a) (b) (c) 2 sensitivity experiments 3 sensitivity experiments 2 + 1 sensitivity experiments Fig. 11. Overflow transport (Sv) ð1sv¼ 10 6 m 3 =sþ of the Red Sea water through the Bab el Mandeb Strait, through a crosssection at point A in Fig. 1d. Details of the estimation are provided in the first paragraph of Section 3.2. The solid line in each panel shows the result of the reference experiment ( refrun ). (a) Two sensitivity experiments ( buoy-run and wind-run ) in Section 4.1. (b) Three sensitivity experiments ( indian-run, aden-run and red-run ) in Section 4.2. (c) Three sensitivity experiments ( arab-run, equa-run and fpln-run ) in Section 4.3. annual mean wind. The heat and water fluxes are monthly over the whole computational domain (as in the reference experiment). For the second and third experiments, the monthly wind is limited to the Gulf of Aden ( aden-run ) and the Indian Ocean ( indian-run ), respectively (the two regions are separated at 501E). Fig. 11b shows that red-run and aden-run produce almost identical results, with less significant seasonal evolution in the overflow rate. Interestingly, however, the monthly wind over the Indian Ocean ( indian-run ) produces a maximum transport of 0.48 Sv in March May (end of the northeast monsoon) and a minimum transport of 0.24 Sv in August (end of the southwest monsoon), which resembles the reference run. These results suggest that the monsoonal wind in the Indian Ocean can remotely control the water exchange through the Bab el Mandeb. To elucidate the dynamic link between the Indian Ocean and the Bab el Mandeb, we have next investigated the depth of the thermocline in the Arabian Sea and the Gulf of Aden. Fig. 12 shows the depth of a 26:0s y isopycnal during the final year of the reference experiment. Its depth is about 200 m in all months except September when the Gulf of Aden shoals in response to the southwest monsoon (cf. Luther and O Brien, 1985; Bauer et al., 1991; Manghnani et al., 1998; Smith et al., 1998; Vecchi et al., 2004). The lifted isopycnal (about 120 m deep) in the Gulf of Aden clearly continues to the Red Sea through the Bab el Mandeb, which leads to blocking of the Red Sea outflow during this period. The hydraulic theory of Siddall et al. (2002) has already explained that the Red Sea outflow minimizes when an isopycnal in the Gulf of Aden is uplifted. The shoaling is consistent with the intrusion of the Gulf of Aden intermediate water to the Red Sea (Section 3.2). The late summer shoaling of the Gulf of Aden is also evident in the World Ocean Atlas. In the monthly climatology, the depth of a 26:0s y isopycnal is about 180 m in February March and about 70 m in August September, which extends over the whole region of the Gulf of Aden. We have produced Fig. 13 which shows time evolution of the depth of upper thermocline, averaged over the western Gulf of Aden, for all experiments. The reference experiment ( ref-run, solid line), wind-run, and indian-run demonstrate a distinct seasonal cycle of 200 m deep in March and 130 m deep in August September. In contrast, without seasonal wind forcing in the Indian Ocean, buoy-run, red-run, and aden-run show little

1462 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 Fig. 12. Seasonal evolution of the upper thermocline depth (shaded, [m]), defined by a 26:0s y isopycnal, during the 15th year of the reference experiment. Climatological wind stress vectors ðn=m 2 Þ are superimposed. seasonal variability in thermocline depth. The shoaling of the western Gulf of Aden is evidently caused by the monsoonal winds over the Indian Ocean, instead of the local wind forcing. Each timing of the deepest depth (March) and shallowest depth (August September) matches that of maximum and minimum outflow in Fig. 11. This result is consistent with Siddall et al. (2002) in that the seasonal cycle of thermocline depth in the Gulf of Aden is in phase with that of the bottom water outflow at the Bab el Mandeb. 4.3. Which part of the Indian monsoon is important? We have conducted two further sensitivity experiments to explore which part of the Indian monsoon is important for the seasonal water exchange through the strait (cf. Bruce et al., 1994; Sengupta et al., 2001; Brandt et al., 2002; Prasad et al., 2005; Rao and Behera, 2005). We have limited the monthly wind stress to the Arabian sea (north of 101N) and then to the equatorial Indian Ocean (south of 101N), with a steady wind stress elsewhere

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1463 depth [m] depth [m] depth [m] (a) (b) (c) 2 sensitivity experiments 3 sensitivity experiments 2 + 1 sensitivity experiments Fig. 13. Same as Fig. 11 except for the depth (m) of a 26.0s y isopycnal, averaged over the western Gulf of Aden, west of 451E. of each. In Fig. 11c, for the case of the seasonal wind over the Arabian sea ( arab-run ), the overflow rate shows a maximum (0.47 Sv) in March and a minimum (0.28 Sv) in August, whereas the second case with variability in the equatorial Indian Ocean ( equa-run ) does not show such enhanced seasonality. Also Fig. 13c shows that the western Gulf of Aden shoals when wind stress changes seasonally over the Arabian sea. These indicate that the monsoonal wind in the northern side of the Indian Ocean is essential for the seasonal water exchange at the Bab el Mandeb, as has been presumed by Jones and Browning (1971). In an attempt to confirm the above conclusion, we present an analytical theory for the coastal upwelling in the Gulf of Aden. We adopt a 1.5-layer planetary geostrophic model on an f-plane: fhv ¼ g 0 hh x þ t x =r 0, (4) fhu ¼ g 0 hh y þ t y =r 0, (5) h t þðhuþ x þðhvþ y ¼ 0, (6) where f is the Coriolis parameter, h and g 0 are layer thickness and reduced gravity appropriate for the definition of a 1.5-layer model, ðu; vþ and ðt x ; t y Þ are the eastward and northward component of velocity and wind stress, respectively. Taking curl of Eqs. (4) and (5) to obtain a form of velocity divergence, we rewrite the continuity Eq. (6): h t ¼ ty x tx y. (7) f r 0 An areal integral of Eq. (7) over the Gulf of Aden yields h t ¼ 1 Z t y j Bf r x¼51 E dy, (8) 0 where B is the surface area of the Gulf, west of 511E where the tip of the Somalian Peninsula locates. The right-hand side of Eq. (8) is derived by assuming that the line integral of wind stress along the northern, western, southern boundaries of the Gulf (Fig. 14a) is negligible. It follows from Eq. (8) that the seasonal shoaling of the thermocline in the Gulf of Aden is controlled by the seasonal variability in the meridional wind in between the Gulf of Aden and the Arabian Sea. The theoretical rate of shoaling and deepening is estimated in Fig. 14b using the wind stress of the reference experiment. If this rate is integrated for a year, the thermocline depth can change seasonally 40 m, which is comparable to the simulated variability in Fig. 13. Also in Fig. 14b, the sign reverses in April May and in September October. Each timing lags behind the period of deepest and shallowest depth in the western Gulf in the reference experiment (Fig. 13)

1464 ARTICLE IN PRESS H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 GULF OF ADEN 51E (a) (b) Fig. 14. (a) Schematic of the route of line integral on the right-hand side of Eq. (8). (b) Theoretical rate of coastal Ekman down/upwelling (positive/negative values) (m/month) over the Gulf of Aden, given by Eq. (8) for the wind stress of the reference experiment with B ¼ 2:35 10 11 m 2, f ¼ 3:0 10 5 s 1, and r 0 ¼ 1026:0kg=m 3. Anomaly from the annual mean is shown. by more than 1 month. This discrepancy suggests importance of other processes neglected in the above simplified theory, such as impacts of subsurface layers, fluxes through the Bab el Mandeb, and nongeostrophic dynamics: these examinations are devoted a later study. We have conducted a final sensitivity experiment ( fpln-run ) by fixing the Coriolis parameter over the Arabian marginal seas (the Arabian sea, the Gulf of Aden, and the Red Sea). This intends to support the f-plane assumption used in the above theory and also to investigate the impact of Rossby waves on the seasonal exchange at the Bab el Mandeb. The result (see Fig. 11c) turns out to be almost the same as the reference run in that there is a winter maximum and a summer minimum, even without the effect of Rossby waves over the northwestern Indian Ocean. We have also checked if any pattern of coastal Kelvin waves is found in the thermocline depth during the period of shoaling in the Gulf of Aden (cf. Shankar and Shetye, 1997). It turns out that the thermocline shoals almost simultaneously over both the Gulf of Aden and the coastal regions of the Arabian sea (not shown). This is

H. Aiki et al. / Continental Shelf Research 26 (2006) 1448 1468 1465 contrasted with the dynamics of subsurface layers: in Section 3.3 we show that patches of IOIW propagate westward along the Arabian coast at the mid-depth. The upper layer and intermediate layer seem to exert different time scales in the Gulf of Aden. The analyses in Sections 4.1 4.3 have presented both numerical and theoretical evidences for explaining how the Indian monsoon regulates remotely the late summer stopping of the Red Sea outflow. The Red Sea outflow is regulated by the seasonal shoaling of the Gulf of Aden as a part of the Indian monsoon system. It may be possible in a future study to deduce the interannual variability of the water exchange at the Bab el Mandeb from reanalysis data of monsoonal wind over the Indian Ocean. 4.4. Discussion Some readers may find that the experimental design of ref-run, wind-run, and buoy-run in the present study is analogous to that in the E1, E4, and E5 experiments, respectively, in Sofianos and Johns (2002). A careful examination is needed for the definition of layer interface, in discussing the twoand three-layer exchange at the Bab el Mandeb. In Sofianos and Johns (2002) layer interfaces are defined by fixed density values, rather than the zero-crossing in velocity profile. Hence both the surface and middle layer transports in their Figs. 7 and 8 are directed to the Red Sea, and have little relevance to the three-layer exchange in observational studies. Moreover, the surface layer in Sofianos and Johns (2002) has densities less than 25:5s y, which include the observed middle layer density of 24:8s y in Table 1. With insufficient vertical resolution for upper layers, isopycnalcoordinate models can produce misleading dynamics of water exchange at the Bab el Mandeb. Nonetheless, comparison with the results of Sofianos and Johns (2002) follows. The seasonal difference in the bottom outflow is 0.4 Sv in ref-run and about 0.2 Sv in the E1 experiment, which may be compared with observed estimates of 0.7 Sv ( ¼ 0.7 0 Sv) in Murray and Johns (1997) and 0.56 Sv ( ¼ 0.61 0.05 Sv) in Sofianos et al. (2002). While the wind-run is successful in reproducing the main characteristics of ref-run, neither of the E4 and E5 experiments reproduces the seasonal bottom outflow of the E1 experiment. The E4 experiment reports reduction in the annual mean bottom transport due to absence of seasonal cycle in the thermohaline forcing, which does not contradict the present result from wind-run. The above-mentioned improvements in the present study over the result of Sofianos and Johns (2002) are brought by finer model resolution (both vertical and horizontal) and improvements in outer boundary condition. The conclusion derived from the present study concerns only impacts of the large-scale wind. The present model used the ECMWF wind which significantly underestimates the local wind over the Bab el Mandeb. We show that, even using the largescale wind, one is able to reproduce a seasonal difference of 0.4 Sv ( ref-run ) in the overflow rate of RSOW, of which at least a seasonal difference of 0.24 Sv ( indian-run ) is caused by the monsoonal winds over the Indian Ocean. Satellite winds will present an important comparison with the result of the present study. Also devoted to a future study is a simulation with increased horizontal resolution up to 0.5 1 km, which will enable a more realistic simulation of the Northern and Southern Channels, especially the nonhydrostatic formulation. Although integration time is limited to less than a few 10 days, several recent studies attempt such high resolution simulation (cf. Legg, 2004; Nakamura and Awaji, 2004). 5. Summary The mechanism of seasonal cycle in the Red Sea water outflow has been investigated in the present study using a locally high-resolution global OGCM, in which the horizontal resolution varied from 2 km around the Bab el Mandeb Strait (the narrow exit of the Red Sea), to less than 20 km in the Gulf of Aden and about 80 km in the Indian Ocean. Although the narrow topography of the Bab el Mandeb leaves several challenges (such as the insufficient resolution of the present model to fully resolve the bottom channels, and the excessive computational load due to the CFL condition), the 15 years reference experiment has provided good simulations of the basic features of the Red Sea outflow system, including the evolution of the RSOW tongue in the Gulf of Aden and the summer minimum and winter maxima in the overflow transport through the Bab el Mandeb. Moreover, we have exampled the three-dimensional structure of the RSOW tongue (Section 3). Interestingly, while the tongue meanders, the discharged RSOW (incoming IOIW) is always characterized by anticyclonic (cyclonic)