The mean flow and variability of the Gulf stream-slopewater system from MICOM

Ocean Modelling 17 (2007) 239–276 www.elsevier.com/locate/ocemod

The mean flow and variability of the Gulf stream-slopewater system from MICOM Angelique C. Haza a,*, Arthur J. Mariano a, Toshio M. Chin a,b, Donald B. Olson a a

Division of Meteorology and Physical Oceanography, Rosenstiel School of Marine and Atmospheric Science, University of Miami, FL 33149, USA b Jet Propulsion Laboratory, Pasadena, CA 91109, USA Received 27 March 2006; received in revised form 6 December 2006; accepted 19 February 2007 Available online 12 March 2007

Abstract The slopewater region is influenced by surface wind-driven, deep buoyancy-driven and shelf currents, whose complex interactions can affect both the northward heat transport and southward return flows. The mean flow and variability of the Gulf Stream-slopewater system are studied using four year outputs of Miami Isopycnic Coordinate Ocean Model (MICOM) realistic high-resolution simulation of the North Atlantic circulation. Special attention is focused on the eastward Slopewater Jet (SJ), a surface current characterized by a mean path coinciding with the strong outcropping temperature front in the slopewater. The water mass, path and transport of the SJ in MICOM are found to be in reasonable agreement with the existing observations. The modeled SJ is associated in part with a GS’ secondary branch, induced by a Taylor column effect of the New England Seamount Chain (NESC) on the upper GS. This upper-ocean-topographic coupling results in a spatial GS bifurcation, and advection of GS waters into the slopewater region shortly downstream of the NESC. An EOF analysis of the pycnocline depth confirms this tendency, as the first mode displays a qualitative dependence of the GS fan-shape streamline dispersion on the strength and intersecting latitude of the incident GS. Additionally, the model displays a strong influence of the Deep Western Boundary Current (DWBC) on the path of the SJ, by acting as a potential vorticity barrier. Important interactions between the two currents are suggested by the statistical EOF of the slopewater column, as in observations. Downstream of the NESC, the SJ transport variability is seasonal in MICOM, due to the north-south annual oscillation of the GS path and mergings with anticyclonic eddies. However, the variability of the SJ velocity profile is dominated (49% eigenvalue) by lateral translations of the current, at a 9-month timescale characteristic of GS meander-intensity variability. South of the Grand Banks, the transport variations of both the SJ and Labrador Current (LC) are captured by the first mode of the upper-slope water (37% of the variability), with predominant timescales corresponding to the upstream variability of the GS, and seasonality of the LC. Both modes are also in reasonable agreement with observations. Published by Elsevier Ltd.

*

Corresponding author. Tel.: +1 786 546 1285. E-mail address: [email protected] (A.C. Haza).

1463-5003/$ - see front matter Published by Elsevier Ltd. doi:10.1016/j.ocemod.2007.02.003

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1. Introduction The slopewater region is a narrow area approximately 250 km wide above the continental slope, between Cape Hatteras and the Grand Banks, bounded by the continental shelf to the north, and the northern edge of the Gulf Stream (GS) to the south. McLellan (1957) calculated that the upper slopewater mass is the result of a mixture of 20–50% of coastal waters, and 80–50% of GS waters. The transition between these three adjacent water masses consists of sharp temperature and salinity gradients which are clearly visible from sea surface temperature data and in hydrographic surveys. One of the characteristics of the strong temperature front marking the transition between Shelf and Slope waters is that it coincides with the path of an eastward current called the Slopewater Jet (SJ), flowing parallel to, and north of the GS in the upper slopewater between 40°N and 43°N (Pickart et al., 1999; Olson, 2001). Although it was described in the literature as a persistent current, the proximity of the GS renders measurements of the SJ difficult when meanders venture more inshore and potentially merge with it. Fuglister and Worthington (1951) and Fuglister (1963) observed multiple filaments of currents in this area, including a distinct and persistent eastward current south of the Laurentian Channel. By assuming mass conservation of the area west of the Grand Banks and north of the Stream, McLellan (1957) estimated the transport of the SJ to be between 10 and 20 Sv. Recently, Pickart et al. (1999, hereafter P99) made much weaker estimates of the mean SJ transport, 2–3 Sv at 55°W and 9 Sv at 50°W, which were derived from geostrophic velocities based on combined hydrographic sections measured during cruises between 1959 and 1995. The slopewater region is also a confluence region for subtropical and subpolar waters: the Labrador Current (LC) and Deep Western Boundary Current (DWBC) both flow southwestward along the continental slope, while the GS and SJ flow northeastward, and the subsequent interactions among these currents account for the complexity of the circulation in this area. P99 found, for instance, that the variability of the SJ at 55°W is mostly dominated by lateral oscillations in phase with the path of the GS, whereas at 50°W, the transport variations of the SJ are in phase with the transport of the LC. Similarly, a strengthening of the SJ coincides with a strengthening of the Denmark Straight Overflow Water (DSOW) and weakening of the Labrador Sea Water (LSW). In light of the recent interest in climate anomalies and their connection with the thermohaline circulation, these limited observations emphasize the importance of improving our knowledge in the dynamics of the circulation in this region. Numerical simulations of the North Atlantic basin have improved greatly during the last two decades, particularly in terms of the separation of the modeled GS from the coast, penetration into the basin interior and realistic transport values. Accurate simulations require state-of-the-art ocean general circulation models with a fine horizontal grid-resolution of at least 8 km. For coarser resolutions, the GS tends to remain too close to the coast and separates at a more northern latitude than the observed separation at Cape Hatteras, resulting in an overlap with the slopewater circulation. The recent availability of adequate North Atlantic basin simulations allows for a more in-depth analysis of the GS-slopewater system. One of the latest North Atlantic basin simulations using the Miami Isopycnic Coordinate Ocean Model (MICOM), with a high horizontal grid resolution of 1/12° and 20 density layers, reproduced a reasonable mean path of the GS, and of the SJ. The path of the SJ is depicted by synthetic drifter trajectories (Haza, 2004) as a persistent current independent of ring activity, while 10 day-averaged surface velocities and temperatures (Fig. 1) display an eastward-flowing current in the slope region, quite distinct from the GS, and evident in all seasons. Associated with this current is a relatively strong surface temperature front, indicating the presence of a barrier that inhibits mixing with the colder shelf waters. Furthermore, the Slopewater Jet in MICOM appears to be fed by the GS via a bifurcation occurring in the vicinity of the New England seamount chain (NESC). McLellan (1957) mentions also that the momentum of the Slope Jet is induced by the GS, but suggests a different mechanism than a bifurcation, such as a frictional drag transferring momentum across the northern edge of the Stream. Yet, his estimate of the location where the GS and SJ are the closest, and where coastal and GS waters are the most likely to meet coincides with the northern tip of the NESC. This raises the question whether an upper ocean-topographic coupling linked to the seamount chain is the main mechanism for the apparent bifurcation of the GS, leading to the emergence of a secondary current in the slope waters. The impact of the NESC on the path of the GS has been previously

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Fig. 1. Ten-day-average surface velocities and temperatures of MICOM outputs during year 1980 in the winter (a) and summer (b). The bottom topography contours are the of 0-, 350- and 4000-m isobaths. Note the presence of the Slopewater Jet along the 4000-m isobath and the temperature front.

investigated in the oceanographic community, although conclusions from observations remain contradictory. Some measurements of the GS path display a distinct change of regime in the amplitude and meandering activity shortly downstream of the NESC (Richardson, 1983; Kelly, 1991). Yet it was argued that local atmospheric forcing could also be the cause of a regime transition in the vicinity of the NESC (Ezer and Mellor, 1992). From 30 months of AVHRR SST-data, Cornillon (1986) found no variation in the rate of change of meander amplitude and envelope of the GS. Lee and Cornillon (1995) later added, using the same method, that only the less energetic meanders are affected by the seamount chain. Other characteristics of the

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NESC-effect on the GS include the observation of a southern deflection of the Stream with a quasi-stationary meander downstream of the chain (Richardson, 1981; Teague and Hallock, 1990), and existence of eddies trapped above the seamounts (Huppert and Bryan, 1976; Richardson, 1981). The southward deflection of the Stream was also reproduced numerically by Adamec (1988) in an idealized configuration using a perpendicular seamount chain, Ezer (1994) with a coastal ocean model, and Blayo et al. (1994) from assimilation of altimetry data into a multilayered quasi-geostrophic model. One of the main mechanisms considered for this upper ocean-seamount coupling is the Taylor column (Vastano and Warren, 1976; Richardson, 1981; Ezer, 1994). Analytical (Hogg, 1973; McCartney, 1975) and numerical (Chapman and Haidvogel, 1992) studies of such effect were conducted only on an isolated seamount, with a horizontally uniform steady flow, a configuration that can hardly be applied to the New England seamount distribution, and time-dependent jet-like profile of the GS. Adamec (1988)’s study of a double-gyre circulation over a perpendicular seamount chain in a 2-layer QG-model is the sole exception. Most of these studies focused on the main topographic effect on the GS, i.e. how it is deflected and how it changes the characteristics of the meanders downstream of the chain. But the bifurcation scenario, in relation to the NESC and dispersive effect on the upper-layer streamlines, has not been explored yet. Given that this region is a confluence zone of several important surface, coastal, and deep currents interacting with one another and topography, and the resulting dynamics play an important role in the overall circulation, four years of a MICOM high-resolution simulation of the North Atlantic are used to investigate the GS-slopewater circulation, with emphasis on the SJ dynamics and the topographic influence of the NESC on the GS. The first-order statistics of the flow field, temperature and salinity, as well as their temporal variability components are analyzed. The primary purpose is to help shed more light on these complex interactions using the high spatial and time resolution of a realistic OGCM simulation. The model limitations are also highlighted using existing observational data. The paper is organized as follows: The model configuration and analysis are explained in Section 2. The main characteristics and variability of the GS-slopewater system are described in Section 3. The GS-bifurcation hypothesis and its link with the NESC is explored in Section 4, followed by a study of the DWBC’s influence on the path of the SJ in Section 5. A discussion/summary is then presented in Section 6. 2. Methodology 2.1. MICOM outputs The model outputs studied in this chapter are derived from a high-resolution numerical simulation of the Atlantic Ocean, using the Miami Isopycnic Coordinate Ocean Model (MICOM) (see Bleck et al., 1992, and Bleck and Chassignet, 1994, for a detailed review). The total computational domain is the North and Equatorial Atlantic Ocean basin from 28°S to 70°N, including the Caribbean Sea, the Gulf of Mexico and the Mediterranean Sea. Our analysis domain is between 30°N to 48°N, and 78°W to 50°W. The bottom topography is derived from a digital terrain dataset with 50 resolution (ETOPO 2.5 quality-controlled data set). The surface forcing consists of the ECMWF (European Center for Medium-range Weather Forecasts) 6hourly wind forcing from the 1979 to 1999 reanalysis, including surface radiation, specific humidity and air temperature, as well as precipitation from COADS (Comprehensive Ocean and Atmospheric Data Set), and a weak relaxation to the surface salinity. Relaxation to observed monthly T and S (Levitus, 1982) is carried out in 3° wide buffer zones with relaxation times of 5 days at the northern and southern wall boundaries increasing to 30 days at the inner edge of the buffer zones. The horizontal grid is defined on a Mercator projection with a resolution of 1/12°  1/12°cos / (where / is the latitude), which corresponds to a mesh size of 6 km on the average. A high horizontal grid resolution is necessary to reproduce an inertial boundary layer, ¨ zgo¨kmen et al., 1997). which is an important factor in the Gulf Stream’s separation from the coast (O The vertical density stratification is composed of 19 isopycnic layers and an active surface mixed layer. The model is spun up by starting from the final state of a previous 20 year simulation forced by an annual climatology. The diapycnal mixing is dependent on the Richardson number. The entrainment parameterization is from Hallberg (2000). The thickness diffusion is 0.3 cm s1  mesh-size. The momentum mixing is composed of a

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Laplacian momentum mixing which is the greater of 0.75 cm s1  mesh-size and 0.1  mesh-size2  velocitydeformation, plus a small biharmonic momentum mixing of 1 cm s1  mesh-size3. The temperature and salinity mixing is 0.5 cm s1  mesh-size, and the diapycnal diffusivity is 0.001  buoyancy-frequency1 cm2 s2. The output fields available for this study are daily averages of the model variables for the surface mixed layer, and 3-day averages for all layers. The description of the mean flow is often simplified here by approximating the 20-layer configuration as a 2 layer model with the pycnocline as the layer-interface, which we assumed to be the interface separating layers 11–12. The resulting mean path of this model configuration is very similar to the mean path of the previous MICOM simulation forced by the monthly climatological ECMWF atmospheric fields, which was found to match reasonably well with observations (Chassignet and Garraffo, 2001). Both simulations include the presence of a slopewater Jet, distinct from a mean Gulf Stream. The highest speeds for the simulated Gulf Stream, between Cape Hatteras and the Grand Banks, range from 30 to 70 cm/s, which correspond to the mean kinetic energy derived by Garraffo et al. (2001) for the archived AOML/NOAA drifter data set, although in the model, there is a stronger decrease in mean velocity intensity in the vicinity of the NESC. However, the eddy speeds in the real ocean (from the in situ drifters) are twice as high as the eddy velocity magnitudes in the MICOM runs forced with both daily and monthly climatological ECMWF, respectively. The low EKE and reduced eastern penetration of the GS relative to observations may be due to the large value of the dissipation operator for momentum. Despite these shortcomings of the numerical simulation, it is one of the most realistic simulations of GS path and structure performed at this time. 2.2. EOF analysis and section choice Vertical and horizontal sections of the domain are selected to study the Slopewater Jet, where EOF-analyzes are performed on the time variabilities of the velocities, temperature and salinity, as well as the pycnocline depth. For example, let v(x, z, t) be the orthogonal velocity of a vertical cross-section at location (x, z). Then v0 ðx; z; tÞ ¼ vðx; v; tÞ  vðx; zÞt is the fluctuation component of the velocity at time t, and vðx; zÞt is the 4-year temporal average of v at that particular location. Each field is thus detrended by the temporal mean with division of the domains into subdomains to zoom-in on a particular current. The three-dimensional domains contain about 104 data points. Since the fields are 9 day averages, the number of time-realizations is 162 over the 4-year study period. The Empirical Orthogonal Functions (EOFs) are then calculated with an algorithm for singular value decomposition, a useful tool for deriving eigenvectors and eigenvalues of data covariance matrices where the number of realizations is equal to or smaller than the number of variables (cf Bretherton et al., 1992). The EOF decomposition for the variable v can be written as: X vðx; z; tÞ ¼ vðx; zÞt þ PCk ðtÞ  EOFk ðx; zÞ; k

where EOFk is the kth (dimension-less) EOF mode, and PCk(t) is the principal component (PC) of mode k at time t, defined as: PCk ðtÞ ¼ Rx;z EOFk ðx; zÞ  v0 ðx; z; tÞ: In most cases for the vertical profiles, the EOF decomposition results in 2–3 dominant modes, with the first two modes explaining 70–80% of the total variability. The interpretation of the modes is facilitated by recombining independently mode 1 or mode 2 with the time Eulerian mean, which helps visualize the average range of variability of each mode. Specifically, adding or subtracting to the mean velocity the normalized EOF modulated by the standard deviation r of its PC time-series displays the probable range of the flow for that particular mode, defined as: t ~v k ðx; zÞ ¼ vðx; zÞ  rðPCk Þ  EOFk ðx; zÞ:

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Note that when two currents such as the SJ and DWBC exhibit much different ranges of variability, the energetic current (i.e. the SJ) dominates the first modes. In order to capture their coupled variability, the correlation matrix is used instead of the covariance matrix to derive the EOFs, where each data is divided by its time standard deviation. As the model outputs constitute a data-set with regular temporal sampling, the dominant timescales of each mode can be extracted by deriving the power spectrum estimate of the PC time-series, using Welch’s averaged periodogram method. This method (cf Welch, 1967) estimates the power-spectra by splitting the data-set into data segments, then calculating the power-spectrum by applying a window function to the time-domain data, computing the discrete FFT and averaging their squared magnitudes. The spectral peaks have a sampling frequency of 0.08 year1, corresponding to a spectral resolution of 1, 0.7, 0.6, and 0.25 months for the relevant periods of 12, 10, 9 and 6 months, respectively.

Fig. 2. (a) Location of all the cross-sections, superimposed on the summer average pycnocline depth contours (cint = 50 m), and the (0, 350 4000)-m isobaths. Average speed of the upper layer is given by the shading. (b) 4-year averaged projected velocities of the vertical cross-sections A, B, C, D (contour interval is 5 cm s1).

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Fig. 3. Time evolution of the projected velocity fields at 50-m depth for the four vertical cross-sections.

The subspace similarity of two EOFs is evaluated by calculating the correlation coefficient qðpcA ; pcB Þ ¼ covðpcA ; pcB Þ=ðrpcA rpcB Þ and considering only values higher than 40%. Due to the high number of degrees of freedom (i.e. 162-2), these correlations are statistically significant. The choice of the sections is dictated both by the need to compare the MICOM outputs with P99’s observations, and by the dynamics of the simulated Slopewater Jet (cf Fig. 2). P99 analyze different quantities from hydrographic measurements centered around 55°W and 50°W. A cross-section at 50.5°W (section C) south of the Grand Banks is defined near the domain boundary. Between the NESC and the Grand Banks, two vertical sections are considered: section D at 55°W for comparison purposes, and the main section closer to the NESC (section A), the latter being more suitable to the EOF analysis of the velocities, because the mean currents in the model are nearly-perpendicular to the zonal axis of this cross-section. The fourth cross-section (section B) is chosen downstream of Cape Hatteras in order to monitor the GS and initial-SJ dynamics upstream of the seamount chain (cf Hovmoller diagram of near-surface velocities in Fig. 3). Additionally, the effect of the NESC on the Gulf Stream-slopewater system is studied by looking at the different modes of variability of the pycnocline-depth. Three horizontal sections are then defined: above, upstream and downstream of the NESC, in order to relate the bifurcation patterns and SJ intensity to the incident path and transport of the Gulf Stream. 3. Gulf stream-slopewater system characteristics 3.1. Mean flow characteristics 3.1.1. Geographic location of the mean path The pycnocline depth contours, calculated by averaging over one year of model output (Fig. 4a) give a reasonable estimate of the geostrophic mean flow in the upper-layer. The flow is more energetic where the

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Fig. 4. (a) Contours of the individual yearly average pycnocline depth in MICOM (contour interval is 50 m), superimposed on the 350and 4000-m isobaths. (b) 4-year averaged velocity contours (in cm s1) superimposed on the gradients of the averaged temperature and salinity at section A. The SJ location corresponds to the surface outcropping of the thermocline and halocline.

contours are more concentrated, and shows the GS separating from Cape Hatteras before the meandering activity starts affecting the coherence of the jet downstream of the NESC. (Note that as the meandering increases east of 70°W, the effects of temporally averaging a meandering current also increase, and this is evident in the mean path losing its strong horizontal coherence.) Yet the GS mean path is still visible, which

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shows clearly the existence of a second and weaker current: the Slopewater Jet, located between the shelf break and the GS’s northern edge: the pycnocline contours concentrate along the 4000–4500-m isobaths. The yearly averages indicate that this current appears to originate from the GS at two different locations: the first branching point is shortly downstream of Cape Hatteras, where the signatures of strong anticyclonic eddies are still visible in the mean-pycnocline, and seem to have deflected a portion of the Stream’s northern edge toward the continental shelf. The second branching point is marked by an anomalous contour dispersion above the NESC, leading to more of the GS being deflected into the slopewater region. Note however, that observations between 74 and 70°W north of the GS depict a southward flowing current along the shelfbreak and slope, merging further with the GS as a persistent recirculation well documented in the literature (i.e. Csanady and Hamilton, 1988), in contrast with the current above the shelfbreak in the MICOM simulation.

3.1.2. Average SJ profile and transport The model SJ referred in this study is defined as the portion of the flow with positive orthogonal velocities above the continental slope, and usually contained in the upper 500 m as it lies above the opposite-flowing DWBC, or separated from the GS by a return flow. The vertical cross-section of the velocities at section A in MICOM (Fig. 2b) displays the model SJ as an upper layer current with a high velocity core extending to a depth of 400 m (similar to P99’s observations), and maximum mean velocities of the order of 20 cm/s. The GS is represented by the other jet profile with a core depth of 800 m and strong barotropic velocities which extends to the bottom. The DWBC corresponds to the area of maximum negative velocities against the continental slope. The transport of the SJ varies greatly depending on the location (Table 1): it increases overall from roughly 3 Sv between Cape Hatteras and the NESC, to 7 Sv shortly downstream of the NESC, and reaches 19 Sv south of the Grand Banks. During the years 1981–1983, the annual transports are 1.5–2.5 Sv upstream of the NESC and 6.5 Sv downstream of the NESC. Comparatively, the 1980 transport is 8.6 Sv upstream of the NESC, due to the occurrence of a strongly kinetic GS’s eddy-shedding in summer, and 10 Sv downstream of the NESC. Near 55°W west of the Grand Banks, the slopewater area has a maximum width, giving more range to the GSmeandering activity, thus affecting the distinction between both currents; the SJ is less structured, with the presence of two cores on each side of the DWBC. Yet the overall transport is higher and varies between 8.6 Sv (in 1981) and 15.3 Sv (in 1983). South of the Grand Banks is a zone of convergence for all the currents, which explains the higher transport of the SJ. Annual transports range between 17.5 Sv and 25.7 Sv (1982). Although the transport values are 2–3 times higher than the measurements of P99, they are similar to McLellan (1957) estimates, and we find as well as P99, that the SJ transport at 50°W is three times higher than the transport shortly downstream of the NESC. The absence of a southward shelf current above Cape Hatteras in MICOM may lead however to an overestimation of the SJ transport: as previous estimates of the shelfbreak front are of the order of 1 Sv (Burrage Table 1 Annual, 4-year transports and mean depth of the Slopewater Jet Year

Upstream NESC Section B

Downstream NESC Section A

55°W Section D

50.5°W Section C

1980

8.56 Sv 309 m 2.25 Sv 310 m 1.62 Sv 197 m 2.59 Sv 228 m 3.32 Sv 224 m

10.06 Sv 276 m 6.54 Sv 205 m 6.72 Sv 247 m 6.48 Sv 319 m 7.64 Sv 245 m

11.43 Sv 345 m 8.61 Sv 228 m 11.42 Sv 304 m 15.28 Sv 373 m 9.46 Sv 296 m

20.32 Sv 349/240 m 17.56 Sv 273/260 m 25.66 Sv 386/310 m 17.63 Sv 555/287 m 18.63 Sv 378/256 m

1981 1982 1983 4 years

Note: Section C is given two mean depths, the latter obtained by considering only the positive velocities in the upper 2000-m depth.

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and Garvine, 1988; Rasmussen et al., 2005), the model could overestimate the SJ-transport by as much as 4 Sv upstream of the NESC. 3.1.3. T, S characteristics The northern edge of the GS can be located by a strong surface temperature front separating the cold shelf and slope waters from the GS’s warmer waters. As the meander activity increases further downstream however, the main surface fronts are located along the shelfbreak and along the path of the SJ. For any vertical cross-section between the NESC and the Grand Banks, the T, S-gradient tongues marking the thermocline and halocline reach the surface with a maximum value coinciding with the SJ’s inshore velocity front, as shown in Fig. 4b. Since the isopycnal mixing occurs along the discrete r of each individual layer in MICOM, a more suitable alternative to the T, S scatter-plot is the r,S diagram, as displayed in Fig. 5: the concentration of salinity values (i.e. number of r, S realizations per salinity range of 0.1 psu) provides a better visual representation of the isopycnal mixing and water-mass proportions present in the SJ. In all cross-sections, the GS water mass (corresponding to the saltiest water-types for layers 2–13) dominates in the positive velocities of the flow above the

LAYER INDICE

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Fig. 5. r, S diagram of the positive projected velocities above the continental slope; (a) between Cape Hatteras and the NESC; (b) downstream of the NESC; (c) at 55°W and (d) south of the Grand Banks, showing the concentration of salinity values within a salinity interval of 0.1 psu versus the MICOM density layers. The colorbar displays on a log-scale the number of r, S values within a salinity interval over the four year simulation period. Note the dominance of the most saline GS waters, and the progressive entrainment and mixing of fresher waters from the shelf and LC. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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continental slope, with 80–90% of GS waters present within each layer. This water mass proportion is consistent with McLellan’s 80% high end estimate of GS waters present in the upper slopewater. Two other water masses are gradually entrained on the eastward flow: the shelf water, represented by the light blue band in layers 5 and 6 (50–100 m depth) with salinities of 33–34 psu, and the LC water (100–400 m depth) in layers 8–12 with salinities of 34–35 psu. The highest concentration of LC water occurs below the Grand Banks (cf Fig. 5d). 3.2. Variability of the slopewater system 3.2.1. Seasonal trends of the slopewater system As the GS has a pronounced seasonality, it is assumed that the same must be true of the SJ. The estimation of the SJ’s variability remains approximate, since this current is not strictly independent from the GS. The northeast transport of the SJ is calculated by including only the positive component of the velocities above the continental slope. The positive transport at section A (downstream of the NESC) tends to vary annually, reaching a maximum value in the fall, and a minimum around the end of winter/beginning of spring (Fig. 6). Although less visible, the transport extrema of the DWBC occur at the same time of the year in the numerical model, as displayed by the monthly trends (Fig. 6e and f). The spectrum of the 4-year time series confirms the presence of a significant annual cycle for this vertical cross-section (not shown). The seasonality of the DWBC between the NESC and 55°W as seen in MICOM, is however not supported by observations: while Schott et al. (2004) find a very weak seasonal cycle for the DWBC at the Grand Banks around 43°N from 2 sets of moorings of 2year time series each, the variability seems to be dominated by the fluctuations and meandering of the North Atlantic Current. The transport seasonality of the SJ is directly related to a feeding mechanism from the GS: the GS shifts laterally inshore from spring to summer, bringing its path closer to the SJ. The proximity of the GS in summer and fall leads to the presence of anti-cyclonic eddies increasing the overall NE transport of the slopewater column, and potential merging with the SJ. This merging is enhanced by the tendency of the simulated GS to be more convoluted in the fall (Halliwell and Mooers, 1983) between the NESC and 55°W, in the form of a detached anticyclonic meander (or ring) merging with the SJ, and a cyclonic meander downstream of Gregg–San Pablo, resulting in a more southern path of the GS (Fig. 6d). Whether the GS water advected by the anticyclonic eddy remains in the slopewater region is not clear; it appears that part of the advected flow remains in the SJ, while another portion returns to the Gulf Stream. Although the NE transport at section A has a clear annual signal, the positive slopewater transports further downstream up to 50°W (not shown) have a 10–11 month signal. The eastward transport at 55°W displays a strong offshore contribution, which indicates further interaction with the GS. The tendency of the simulated GS meandering activity to spread over the slopewater region in the summer and in the fall can also be seen in the seasonal T, S characteristics. The GS water mass is recognizable on a T, S diagram as the curve connecting the saltiest water-types within each rh for the upper and intermediate depths. In this MICOM configuration, the relevant rh values correspond to the discrete density layers 2–12 (the mixedlayer is not included). The presence of the GS in the slopewater region corresponds therefore to a high concentration of warmer, saltier (T, S) values along that curve, and relatively lower concentration of fresher, less saline Shelf and LC waters. The two main water types were defined as the max (T, S) values (GS water) and min (T, S) values (Shelf and LC water) of each layer. Following the method used by Garraffo et al. (2003), the water-mass proportions within each layer were calculated based on the assumption that the mixing is mostly isopycnal, and function of the relative distance of each (T, S) dot with respect to the water types of same rh. The overall water mass proportion of the slopewater column downstream of the NESC for the fall and spring seasons over the 4 years of simulation (Fig. 7a) shows a higher concentration of GS waters during the fall for the layer indices 2–9, corresponding to the upper circulation. In spring, there appears to be a higher concentration of the cooler/fresher water masses of the shelf and LC waters. This behavior is highlighted by the salinity concentrations within each layer for the positive projected velocities of each individual season (cf Fig. 7b): the salinity distributions of layers 5–8 show again the stronger mixing of the GS with the fresher water types in spring, versus the purer GS waters in the SJ in the fall.

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Fig. 6. Seasonal averages of the transverse velocities at section A for winter (a), spring (b), summer (c) and fall (d). Contour units are in cm s1. Below are the 4-year average transport values for each month (thick line) of the positive velocities above the continental slope (e), and the DWBC (f) downstream of the NESC. The dashed lines are the monthly standard deviations. Note: the curve is repeated over the months 13–24 in order to help visualize the transport seasonality, including the addition of the standard deviation.

Near and south of the Grand Banks, the salinity characteristics of the SJ display a seasonality due to its interaction with the LC (Fig. 8a). The retroflection of the LC indeed contributes to the transport increase of the SJ via entrainment processes: the concentrations of salinity values versus isopycnic layer for the eastward velocities (Fig. 8b) show a fresh-water branch quite distinct from the GS water branch, corresponding

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100

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% SALTIEST WATER TYPE

70

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LAYER INDICE

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spr.1981

sum.1981

fal.1981

win.1982

spr.1982

sum.1982

fal.1982

win.1983

spr.1983

sum.1983

fal.1983

34 36 SALINITY (PSU)

34 36 SALINITY (PSU)

34 36 SALINITY (PSU)

34 36 SALINITY (PSU)

5 10 15

5 10 15

5 10 15

5 10 15

0

5

10

50

100

500 1000

Fig. 7. Percentage of saltiest water type (a) within each isopycnal layer for spring and fall at section A. The salinity values are split into 2 categories corresponding to the sign of the projected velocities. (b) Seasonal salinity-scatter plot displaying the concentration of salinity values within a salinity interval of 0.1 psu above the continental slope versus the MICOM density layers for the positive projected velocities.

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5 10 15

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Fig. 8. Same as Fig. 7 at section C.

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to the LC between layers 8 and 12. The salinity concentrations are lower in summer and larger in the other seasons, with high values in spring for layers 8 and 9. 3.2.2. EOFs analysis of sections A and C The variability of the GS-slopewater system is now studied by an EOF analysis of the orthogonal velocities and T, S characteristics of the two main vertical cross-sections, downstream of the NESC (section A), and at 50°W (section C). 3.2.2.1. Velocity, T, S EOFs at section A. The section is reduced alternatively to fit to a single current by zooming on its averaged location (cf Fig. 9). As displayed in Table 2, the first two modes of the GS (47%, 33%) and SJ (49%, 24%) explain most of their domain’s respective variability. These modes belong to either one of two main variability subspaces, based on the spectra of their principal component (PC) time series, as explained below: GS-mode 2 and SJ-mode 1 (hereafter GS-2, SJ-1) refer to the same variability subspace, as the correlation coefficient of their PC time-series q(GS-2,SJ-1) = 75%, which is characterized by the common dominant periods of 8–9 months and 3–4 months (Fig. 11b and c). The probable range of SJ-1 in Fig. 9c shows how the velocity-core of the SJ oscillates laterally with little intensity change between its more inshore position adjacent to the DWBC (marked by the zero-velocity contour) and its offshore position at 360 km, which corresponds actually to the northern edge of the GS (identified by the deeper velocity core of the GS at 600–700 m). The ‘‘inshore” GS position in the probable range of GS-2 in Fig. 9b indeed contributes to the SJ variability, resulting in a standard lateral oscillation range of about 100 km for the SJ. Therefore in the GS-2/SJ-1 subspace, the GS feeds the SJ by way of its lateral translation resulting in a partial merging of the two currents with a strong periodicity of 8–9 months. Based on previous analyzes of the GS path from AVHRR infrared images, Lee and Cornillon (1995) extracted a similar dominant period of 9 months for the temporal variation of meandering intensity. The Stream’s lateral displacement captured by GS-2 must then be related to its meandering activity. Note that there is not much change in intensity of the GS in that subspace, which rules out seasonal variations in transport. SJ-2 and GS-1 belong to the second variability subspace, with q(GS-1,SJ-2) = 53%, characterized by the common dominant periods of 1 year and 6–7 months (cf Fig. 11a and c). The probable range of GS-1 (Fig. 9a) confirms the seasonal trends previously discussed. The convoluted behavior of the GS towards the end of summer corresponds to the strong inshore GS-signature in GS-1, followed by the translation of an anticyclonic eddy or high amplitude meander into the slope waters around fall. The meander/eddy-signature appears as the velocity core-dipole on the east side of the offshore GS (Fig. 9b). The merging of the meander/eddy with the SJ results in an increase in momentum and eastward transport of the SJ as seen in the SJ-2 modal state (Fig. 9d), thereby weakening the GS’s main branch. The first two modes for the temperature and salinity of the upper slopewater refer to the same subspace, and their spectra show a strong annual signal in Fig. 11e and f. Mode 1 explains 35% and 32% of the total variability for T and S, respectively, and displays a variation in intensity and lateral shift of the surface fronts (Fig. 10a and c). While the lateral shift of the T, S-front is also of the order of 100 km, this seasonal variability does not correspond to the PC time series of the velocity modes GS-2 and SJ-1. The T, S variabilities at depths deeper than 100–200 m, however, follow the variabilities of the GS and SJ velocities. Therefore in this region, the seasonality of the mixed layer in MICOM dominates the T, S-variabilities for the upper-slope water domain, and does not provide additional information on the GS-SJ dynamics (unless the mixed-layer is truncated from the EOF-domain). 3.2.2.2. Velocity, T, S EOFs at section C. South of the Grand Banks, the Labrador Current flowing in the opposite direction between the continental shelf and the SJ adds another degree of freedom in the variability of the slopewater system. The EOFs of a subdomain named ‘‘upper” are therefore calculated, similar to the observations of P99 (as the only available reference), and composed of both the LC and SJ in the upper 1200 m. The velocity EOFs of each current including the GS are also calculated in order to isolate their individual contribution (cf Table 4). The first two modes of the upper-slopewater explain 37% and 18% of the total variability. The probable range of upper-1 in Fig. 12d captures a variation in velocity magnitude of both the

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Fig. 9. The probable range for each of the first two velocity modes of the Gulf Stream (a, b) and the SJ (c, d) at section A (rPC is the standard deviation of the principal component time-series). Contour interval is 5 cm s1. The thick line marks the zero-velocity contour.

SJ (positive velocities) and the LC (negative velocities), with periods exceeding a year, 1 year, and 10 months (cf Fig. 14a). The transports of both currents are increasing and decreasing in phase, from (70,20) cm s1 to (35, 10) cm s1 for the max SJ and LC velocities, respectively (Fig. 12d). This synchronized transport variability is consistent with the results of P99. By looking at the SJ and LC individually, one can see in Fig. 12b and c, a transport variability in both cases captured by a dominant mode (43% for the SJ, 68% for the LC). The PC-correlations between upper-1 and

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Table 2 Eigenvalues and dominant periods of the principal component time-series of the first 2 EOFs for section A Section

Area/current

Variable

k1 (%)

PC1-periods (months)

k2 (%)

PC2-periods (months)

A A A

Gulf Stream Slopewater Jet DWBC

Vel Vel Vel

47.4 49.1 32.3

12, 6.7 8.6, 3.3 >12, 12

33 24 22.3

8.6, 3.3 >12, 12, 6.7 7.5, 6

A

Upper layer

A

DWBC

A

Slope column

T S T S Vel S

35.3 32.2 48.1 70.2 30 56

12, 8.6 12, 6 >12, 12 >12, 6, 12 12, >12, 5 >12, 8.6, 12

24.5 18.7 14.4 10.7 22 7.6

>12, 12 12, 8.6 >12 >12 >12, 10, 4.6 >12, 10, 4.2

Values of the EOFs of the Slope-column are derived from the correlation matrix. ‘‘>12” denotes dominant periods longer than 12 months, whose values cannot be ascertained with any statistical significance from the 4 year time-series.

each individual current yield however a dominance of the SJ variability in the upper-slope domain with a 98% value (versus 48% for the upper/LC correlation). The power spectrum estimates of the PC time-series in Fig. 14 indicate that the intensity variability of the SJ has predominant periods of >1 year and 10 months (Fig. 14c). As upper-1 and SJ-1 refer to nearly-identical subspaces, it is surprising that the 1-year period does not appear in the SJ’s main variability, and is captured instead by the second SJ-mode (cf Table 4). The only contribution to the annual signal in upper-1 seems to come from the LC-1 variability, which has a pronounced seasonality (cf Fig. 14d). On the other hand, the 10-month signal appears to be the intrinsic frequency of SJ-1 south of the Grand Banks. The 10-month signal may be related to the first velocity mode of the SJ at the upstream section A, i.e. to the meandering intensity variability of the GS downstream of the NESC that is characterized by a similar 9-month period. The presence of the 10-month-dominant signal in the SJ-1 velocity mode at the intermediate section D (cf Table 3) seems to fit with this assumption. The first two upper slopewater T, S-modes explain (44%, 20%) of the temperature’s variability, and (31%, 13%) of the salinity’s variability. Here again, the PC correlation of 96% between T  1 and S  1 implies an identical subspace. The probable range (Fig. 13a and c) displays not much change in the frontal structure for the first mode, and most of the variability is related to the T, S-intensities of the warm/salty tongue, corresponding to the SJ, with a dominant period of 10-months. The annual signal is not present in the first mode, contrary to the upstream cross-sections. This implies that the mixed-layer seasonality does not dominate the variability of the upper 1000 m slope-region at 50°W. The static positions of both the SJ and LC (due to the topographic constraint of the Grand Banks) contribute to maintain the surface T, S-fronts at approximately the same latitude, with relatively small lateral oscillations. These oscillations however, correspond to the PC time series of T  1, S  1, as well as the velocity mode SJ-1, the latter capturing mostly the transport variability of the SJ. It can be inferred that T, S  1 relates indirectly to the transport variation of the SJ, which affects the T, S variabilities by slightly shifting the T, S front in the upper 300 m, as well as the T, S maxima of the warm/salty tongue, which is confirmed by the dominant 10-month period. The second salinity mode is biased by numerical discontinuities near the shelf, as seen in Fig. 13d. However, T  2 clearly captures the seasonality of the upper-slope water (Fig. 13b), with a dominant annual signal (Fig. 14g). Therefore, the dominant mode captures essentially the SJ variability with a 10-month period, while the annual signal, imposed by the LC and the mixed-layer seasonality, is represented by the second mode. 4. NESC influence on the Gulf Stream 4.1. Effect in the mean flow 4.1.1. Taylor column The divergence of the pycnocline contours in the vicinity of the NESC (cf Fig. 4a) raises the question of a possible upper-ocean-topographic coupling that would affect the path of the GS. The NESC is composed of 8

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Fig. 10. Probable range of the first two modes for the mean upper slopewater temperature (a, b) and salinity (c, d) sections downstream of the seamount chain (section A).

seamounts with a base of approximately 50 km, and a top reaching minimum depths of 1600 m (Kelvin) up to 904 m (Gregg). In the MICOM configuration studied here, the horizontal grid-resolution (6–8 km) is not sufficient to capture the peak of the seamounts, thus reducing their minimum depth to a range of 2000– 2500 m (cf Fig. 15). The main influence of the seamount chain captured by the numerical model should therefore be analyzed through potential vorticity considerations. The impact of the NESC on the GS flow is well captured by maps of 4-year mean (MKE) and eddy (EKE) kinetic energy in Fig. 16, displaying in both horizontal and vertical sections an abrupt loss of momentum in

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1

2

200 100

150 100 50 0

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12 mon

POWER SPECTRUM

150

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)

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3

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)

Fig. 11. Power spectra of the principal component time-series at section A: (a) GS mode 1; (b) GS mode 2; (c) SJ mode 1; (d) SJ mode 2; (e) upper slope temperature mode 1 and (f) upper slope salinity mode 1.

Table 3 Same as Table 2 for section D (55°W) Section

Area/current

D D

Gulf Stream Slopewater Jet

D

Upper layer

D

DWBC

D

Slope column

Variable

T S T S

k1 (%)

PC1-periods (months)

k2 (%)

PC2-periods (months)

39.7 23.2

3.5, >12, 8.6 10, 12, >12

26.1 18.2

12, 7.5 >12, 5, 7.5

56.5 48.3 54 65 59.2

12 12 >12, 10 >12, 10 >12, 10, 3.7

10.7 14.2 13.5 10.6 8.5

>12, 12, 4.6 12 8.6, 12, 5 5, 7.5, 6 12, 7.5, 5

the vicinity of the seamount chain affecting the coherence of the GS. Note: an upper (lower) layer field is defined here as the weighted vertical average of the field in all layers above (below) the pycnocline, which corresponds to a two-layer representation of the flow. The increase in eddy activity both upstream and

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Table 4 Same as Table 2 for the section C (50.5°W) Section

Area/current

Variable

k1 (%)

PC1-periods (months)

k2 (%)

PC2-periods (months)

C

Gulf Stream Slopewater Jet Labrador Current Upper layer Upper layer

Vel

44.3 43.2 68.2 36.7 43.9 31.1

>12, 5.4 >12, 10 12, 5.4 12, 10, >12 10 10

25.6 21.7 10.6 18.3 20 13.2

12, 5.4 8.6, 6, 12 6, 8.6 >12, 12, 8.6 12 >12, 12

C

T S

downstream of the NESC indicates lateral excursions of the GS, due to flow obstruction by the seamount chain (Fig. 16c and d). Adjacent to the two farthest seamounts (Gregg and San Pablo), the jet appears constrained to meander anticyclonically, while there is very little kinetic energy between the groups Gregg–San Pablo–Manning and Atlantis II–Gosnold. In the lower layer (Fig. 16b), the aligned chain composed of the first three upstream seamounts (Kelvin, Atlantis II and Gosnold) obstructs for the most part the northeastward flow. Downstream of Gregg and San Pablo, the deep signature of the anticyclonic meander observed in the upper layer (Fig. 16a and c) is visible as the area of highest mean and eddy kinetic energy past the seamount chain. Vertical along stream cross-sections of the MKE and EKE (Fig. 16e and f) show indeed minimum values for the water columns above the seamounts intersecting the GS’s path. The effect is drastic in the lower layer up to the pycnocline, and remains significant in the upper-layer, although slightly weaker. These statistical results hint qualitatively to a possible Taylor Column effect, as the flow below the pycnocline is relatively weak and appears to avoid the seamounts by circulating around them (see Greenspan, 1969, for the unstratified Taylor column and Ezer, 1994, for a scaling analysis applied to the NESC). It remains difficult however to quantify this effect, as most theories apply to idealistic configurations, where a single seamount is intersected by a steady state flow with an idealized density profile. For instance, in the case of a horizontally uniform flow with vertical shear and stratification N intersecting a low circular bump, Hogg (1973) found that the vertical influence is function of a stratification measure S = (Nd/f)2 (where f is the Coriolis parameter, d = H/L, H being the total depth and L the radius of the bump), a topographic parameter b = h0fLu0(0)/U2 (where u0(0) is the bottom velocity, U the average velocity and h0 the height of the bump), and the upstream velocity shear u00 ðzÞ. If S 6  (Rossby number) and b > 2, then the result is a pure Taylor column, whereas S  1 (corresponding to a moderate stratification) leads to a conical vortex. Assuming a steady flow intersecting a New England seamount with approximate linear stratification and constant vertical shear below the pycnocline, the typical scales would be: L = 35 km, h0 = 2500 m, H = 5000 m, U = 15 cm s1, u0(0)  5 cm s1, and   0.04, leading to S  2 and b  1. Hence a stratified Taylor column of conical shape is likely to form above and around the seamount. Similarly, Chapman and Haidvogel (1992) found that true Taylor caps can form in the case of a steady uniform flow on an f-plane with stratification intersecting a tall seamount, if  < 0.15 and h0/H P 0.4. Therefore, a steady uniform flow with velocities below 50 cm s1 intersecting a New England seamount is sufficient to generate a Taylor cap. In spite of the highly time-dependent jet-like profile of the GS, the existence of Taylor Columns related to the New England seamounts has been mentioned in the literature by Hogg (1973), McCartney (1975) and Richardson (1981). It is also worth noting that the fan-shape streamline dispersion of the b-plane inertial Taylor Column derived analytically by McCartney (1975) has some similarities with the GS streamline divergence of the MICOM mean flow above the NESC (cf Fig. 4). 4.1.2. Deep circulation Special attention is thus focused on the deep recirculation cells: The annual lower-layer mean flow (thickness weighted average of the layers 12–17 velocities) in this simulation is quite steady and composed of several recirculation features with scales of a few 100 km (Fig. 17): a cyclonic cell persists upstream and adjacent to Kelvin, Atlantis II and Gosnold. This cell moves slightly northward during certain years, crossing the chain’s barrier by circulating between Kelvin and Atlantis II. Coupled to this cell further upstream is a strong anticyclonic cell visible in all years. Another important feature is the very persistent anticyclonic cell around the Gregg–San Pablo–Manning seamounts.

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Fig. 12. Probable range of the first velocity mode (contour labels in cm s1) at 50.5°W (section C) for areas centered on (a) the Gulf Stream; (b) the Slopewater Jet; (c) the Labrador Current and (d) the upper slopewater.

The combination of quasi-steady recirculation cells with a diameter on the order of 200–300 km has been reproduced numerically with a sigma coordinate ocean model by Ezer (1994), who proved that the presence of the NESC is a necessary condition to the generation of this circulation pattern. Qiu (1994) also found recirculation cells at a depth of 3000 m, derived from combined hydrographic and satellite altimetric data from Geosat. The same recirculation cells are captured, although the cell location does not match perfectly. While

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Fig. 13. Probable range of the first two modes for the mean upper slopewater temperature (a, b) and salinity (c, d) sections at 50.5°W (section C).

Qiu’s determination of the deep flow is indirect, it is based on ocean measurements, thus confirming the existence of such deep recirculation cells as real oceanic features. Because most of the deep flow’s velocities are relatively weak (with respect to the surface flow), it can be explained by a remarkably simple potential vorticity consideration: assume a two-layer configuration with a deep, mostly barotropic flow confined beneath the upper layer, distorted by the GS and the bathymetry. In the absence of forcing or friction, (n + f)/h2 = Constant along the streamlines, where h2 is the lower layer thickness, f if the Coriolis parameter and n is the relative vorticity. The recirculation cells having a length-scale L  200 km and velocities of U  10 cm s1, the relative vorticity n  U/L  106 s1, which is negligible with

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Fig. 14. Power spectra of the relevant principal component time-series at section C: first velocity modes of (a) upper slopewater column; (b) Gulf Stream; (c) Slopewater Jet; (d) Labrador Current, first temperature (e) and salinity (f) modes, and second temperature (g) mode of the upper slopewater column. Note: the second salinity mode is biased by a numerical discontinuity near the shelf and is therefore not represented here.

respect to f  104 s1. Therefore, the potential vorticity in the lower layer can be approximated by the ‘‘potential thickness” (h2/f). The superposition of the lower-layer mean velocities to the lower-layer isopachs of the MICOM output fields illustrates the mechanisms of the deep circulation (Fig. 18): the tendency of the deep flow to follow the h2/f-contours is remarkable, and it indicates that in the mean, the circulation below the pycnocline approximately conserves its potential vorticity. This behavior is particularly clear upstream of the NESC, where the GS is coherent and has a strong steady component. The h2/f-contours seem to fit to the majority of the superimposed recirculation cells, such as: the cyclonic eddy between Balanus and the

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Fig. 15. New England seamount bathymetry used in the high resolution MICOM simulations and derived from the ETOPO 2.5 qualitycontrolled data set. Contour interval is 500 m.

Kelvin subgroup, the two other cyclonic cells further south contributing to the separation of the DWBC from the continental slope, and the anticyclonic eddy circling the Gregg subgroup. The strong anticyclonic subgyre between the deflected DWBC and the cyclonic cell south of Kelvin–Atlantis II is actually the main exception, as it corresponds to a confluence zone where the relative vorticity is not negligible. In the region downstream of the NESC, the flow is significantly weaker partly due to the higher meandering activity, and is therefore less dependent on the isopachs, except for the DWBC flowing along the continental slope. The lower-layer recirculation cells induced by the potential vorticity are enhanced by the presence of the DWBC which flows southward along the continental slope and the lower layer isopachs and, therefore, entrains and is entrained by the deep Gulf Stream flowing toward the northeast. The deep flow interaction between the GS and the DWBC was reproduced numerically by Thompson and Schmitz (1989), who showed indeed that representing the DWBC in their two-layer experiment led to the presence of a northern deep cyclonic gyre between the Grand Banks of Newfoundland and the NESC. This interaction does not only occur when the GS crosses the path of the DWBC after separating from the coast, but also in the vicinity of the NESC, which acts as a barrier and deflects a portion of the DWBC from the continental slope north of the chain.

4.1.3. Downstream topographic effect on the mean Gulf Stream The Taylor column theory, which was invoked as a possible interpretation for floats trapped into eddies above the New England seamounts, stipulates that a steady, barotropic, uniform flow circulates and flows around an isolated seamount anticyclonically, because the conservation of a quantity related to the potential vorticity compensates the bump-induced change of depth by generation of negative vorticity (cf Hogg, 1973). What happens however, in the unknown case of two (or more) nearby seamounts? Whether the flow goes around the two seamounts anticyclonically, or just crosses in between seems to depend on the distance separating them, as one can see from the model. If the seamounts are close to each other, the flow tends to feel the seamounts as a single topographic obstacle of larger horizontal extent and circulates around the seamount cluster anticyclonically, i.e. to the north in the case of an eastward flow. This circulation feature is reproduced in the numerical experiment of Adamec (1988): the deep flow is mostly blocked by the chain of aligned and equally spaced seamounts, and the overall effect on the upper-layer circulation is similar to that of a ridge. There are, however, no recirculation cells present in the lower-layer. It seems therefore, that while Adamec’s configuration is an improvement toward a more realistic representation of the GS-NESC system, the following peculiarities of the NESC were not represented:

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Fig. 16. Four-year mean and eddy kinetic energy of the ‘‘upper” and ‘‘lower” layers, superimposed on the 0-, 250- and 4000-m isobaths. Contour interval is 4 m3 s2 for plot (a) and 2 m3 s2 for plots (b), (c) and (d). Square root of the 4-year mean (e) and eddy (f) kinetic energy of a vertical cross-section intersecting the seamounts Atlantis II and San Pablo. Contour intervals are 5 cm s1 for the 0–800 m depth, 2.5 cm s1 for the 800–1000 m depth, and 2 cm s1 for the 1000–5000-m depth.

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Fig. 17. Annual velocity vectors of the ‘‘lower” layer flow for each year, showing the quasi-steady recirculation cells, and superimposed on the 0-, 350- and 4000-m isobaths.

 The main axis of the NESC is not perpendicular to the mean path of the GS, hence the symmetry of the double-gyre circulation is broken, and the jet’s streamlines encounter the seamounts at different longitudes.  The New England seamounts are not ‘‘strictly” aligned. The relevant ones can actually be split into two groups or subchains: Kelvin, Atlantis II and Gosnold, forming a nearly-perfect alignment, and the second group composed of Gregg, San Pablo and Manning, which are positioned along a curve northeast of the first group’s axis. This configuration can be simplified to an upper-layer jet crossing sequentially over two small seamount chains.  The seamounts are not equally spaced, thus affecting the preferred path of the flow: there is a narrow pass between Kelvin and Atlantis II, while Atlantis II and Gosnold are connected above the ocean floor level, similarly to Gregg and San Pablo. On the other hand, the pass between Gosnold and Manning is relatively wide. One of the main differences between the QG-experiment of Adamec and the numerical simulations using the actual NESC distribution (e.g. MICOM, Ezer experiment) is the downstream effect on the upper layer flow: the perpendicular and linear seamount chain used by Adamec generates a strong anticyclonic meander in the upper layer, followed by a spatial oscillation of the jet downstream of the chain. Similar QG-experiments (cf Haza, 2004) confirm that tilting the axis of the seamount chain affects the symmetry of the lower-layer flow and dramatically reduces the amplitude of the southward deflection. The presence of the seamounts in MICOM simulation triggers two important circulation patterns downstream of the NESC: the upper-layer streamline divergence, which deflects a northern portion of the GS into the slope waters, and a (subsequent) regrouping of the streamlines, contouring Gregg and San Pablo. It seems that the first seamount-cluster interrupts the momentum of the flow and disperses the streamlines, while the second

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Fig. 18. Velocity vectors of the 1980–1983 deep mean flow, superimposed on the lower layer isopachs, or contours of potential-thickness, which is defined as h2  f0/(f0 + by), with a contour interval of 100 m. The solid lines delineate the coast and shelf break.

seamount-cluster adds momentum to the stream by imposing a Taylor constraint (via potential vorticity conservation) on the streamlines, concentrating the flow around Gregg–San Pablo. The effect of the second seamount subchain on the upper-layer flow is comparable to the anticyclonic meander triggered by the seamount chain in Adamec’s experiment, though somewhat weaker. It is, however, the combined effect of these two subchains of seamounts which leads to the reorientation of the stream in two different directions, thereby setting the stage for a GS bifurcation. 4.2. EOF analysis of the NESC influence The impact of the NESC on the GS-SJ variability is studied from several EOF analyzes. The initial stage of the SJ’s variability must then be analyzed and compared with the GS-SJ dynamics above and downstream of the seamount chain. 4.2.1. EOF Between Cape Hatteras and the NESC in the model The projected velocities at section B (vertical section upstream of the NESC) are investigated in that context. The time-evolution of the GS-path displayed in the surface Hovmoeller plot (Fig. 3) shows two remarkable behaviors: the GS’s latitude remains more or less constant between 1980.5 and 1982+, then it is followed by an abrupt shift to a higher latitude before 1982.5 till the end of 1983. There is also episodic anticyclonic eddy-shedding on the upper continental slope (about 3 times during those 4 years), the most energetic event occurring around 1980.5. These two characteristics of the flow help in the analysis of the variability not only at section B, but also further downstream. The dominant mode of section B (Table 5) explains 50% of the total variability, which captures the lateral shift of the GS (Fig. 19a). The corresponding PC time-series (Fig. 19c) shows indeed the abrupt change in latitude between 1982 and 1983. The dominant period is obviously a >1-year period corresponding to the ‘square wave’ generated by the 1982.5 high amplitude latitude shift, and to a lesser extent, the 8–9-month period. The SJ’s dominant mode captures mostly the transport variability of the current adjacent to the continental shelf, and explains 71% of its total variability. The PC time series displays the high amplitude variability due to

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Table 5 Same as Table 2 for the section B between Cape Hatteras and the NESC Section

Area/current

Variable

k1 (%)

PC1-periods (months)

k2 (%)

PC2-periods (months)

B

Full section Slope Jet

Vel

49.9 71.1

>12, 5, 8.6 >12, 8.6

14.8 15.3

5, >12 >12, 12

Fig. 19. Probable range of the first velocity mode at section B, for the full cross-section (a) and the current above the slope, or early SJ (b). (c) and (d) are the corresponding PC time series.

GS eddy-shedding, with an extrema at 1980.5 (Fig. 19d), and dominant periods of >1 year (main eddy shedding), 8–9 months (GS meander-related mode), and 6 months. The EOF-calculations for this cross-section clearly relate the existence of the initial SJ to the GS’s eddy activity between Cape Hatteras and the NESC in the numerical model. 4.2.2. Horizontal EOFs of the pycnocline’s depth The changes in the pycnocline-depth over horizontal portions of the North Atlantic are now investigated to find out how the strength and latitude of the incident GS affect the streamline dispersion above the NESC and

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Fig. 20. PC time series of the first two EOFs of the pycnocline depth for the ‘‘upstream” (a), ‘‘NESC” (b), and ‘‘downstream” (c) domains.

characteristics of the downstream SJ. Due to the energetics of the eddy-field, there are more degrees of freedom on the horizontal plane and the first modes do not capture a significant part of the variability. The region in the vicinity of the NESC is thus split into three areas (about 500 km  700 km) ‘‘upstream”, ‘‘above” and ‘‘downstream” of the NESC (Fig. 2a), in order to obtain higher eigenvalues reaching about 1/5th of the total variability. In the ‘‘upstream” domain however, the main variability of the pycnocline is represented by the first mode with a high eigenvalue of 41%. Although it is a horizontal domain, the presence of a dominant mode is due to the strong coherence of the Stream and low meandering activity. As shown by the PC time series in Fig. 20a and the symmetric probable ranges in Fig. 21a and b, this mode mainly captures the GS lateral shift (and to a lesser degree the GS transport). Note that the trace of a ring above the continental slope is more defined in the case where the GS path is located further offshore. The 95% correlation between this mode and the velocity EOF-1 of section B indicates a quasi-identical subspace, with a common long period corresponding to the 1982.5 latitude jump of the GS. In the area surrounding the NESC, the first 2 modes (21%, 13.5%) display 2 different types of streamline dispersion, although the main focus here is on the first mode. The probable range for mode 1 (Fig. 21) shows the GS oscillating from two states: 1 – a strong incident jet crossing the seamount chain between Kelvin and Atlantis II, followed by a fan shape dispersion occurring shortly downstream (Fig. 21b); 2 – a weaker jet with an upstream meander resulting in a lower angle of incidence, and a loose downstream anticyclonic meander advecting a higher portion of the Stream into the slopewater area (Fig. 21a). The correlation between this mode and the first velocity mode of section B is surprisingly high, with a PC correlation-coefficient of 51%. Both PC-time series share the common periods of >1 year (1982.5 GS latitude jump) and 8–9 months (cf Table 6). The main factor contributing to this mode-similarity appears to be the low-frequency signal present in both PCs, as displayed in Fig. 22a and c, and corresponding to the position of the GS-path. It seems therefore, that the fan-shape dispersion case with strong incident jet of the ‘‘NESC” mode 1 is related to the high-latitude path of the GS, occurring in the years 1982.5+ through 1984, and is depicted in Fig. 21b (The subspace similarity exists also between ‘‘NESC” 1 and ‘‘upstream” 1 with a correlation of 57% as the latter captures the low-frequency signal of the GS latitude). In addition to the common presence of the long square-wave period, the mesoscale frequencies corresponding to 8–9 and 5–6-month periods are highly amplified in the area above the NESC. This analysis suggests that the NESC bathymetry may enhance the GS meandering activity (i.e. 9 month1 frequency), as shown in Fig. 22.

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Fig. 21. Four types of configurations described by the individual modes of the pycnocline-depth, when their PC time-series are more than 40% correlated: a low latitude incident GS and the corresponding flow pattern above and downstream of the NESC (a), versus an incident GS intersecting the NESC at a higher latitude (b). (c) and (d) configurations correspond to the second pycnocline mode of the ‘‘NESC” domain. Pycnocline contour interval is 50 m.

Table 6 Same as Table 2 for the pycnocline depth Section

Variable

k1 (%)

PC1-periods (months)

k2 (%)

PC2-periods (months)

‘‘Upstream” ‘‘NESC” ‘‘Downstream”

Pycnocline depth

40.7 20.8 21.3

>12 >12, 8, 12, 5 5–6, 12, 9

13.2 13.5 19.1

>12, 7.5, 5.2 8, >12, 12 6–7, 5, 12

Downstream of the seamount chain, the first 2 modes explain 21% and 19% of the total variability, respectively, with the proximity of the GS also masking the variability of the SJ. There are significant correlations between ‘‘downstream-1” and ‘‘NESC-2” (71%), and between ‘‘downstream-2” and ‘‘NESC-1” (42%), making

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Fig. 22. Illustration of the influence of the NESC and latitude of the incident Gulf Stream on the pycnocline contours: Principal Component (PC) time series of (a) first velocity mode of the vertical cross-section upstream the NESC; (b) first pycnocline mode upstream; (c) above; (d) and second pycnocline mode downstream the NESC, respectively. (e) Corresponding power spectra.

sense qualitatively from the probable ranges: the shape of the dispersion above the NESC dictates basically the amount of GS water advected in the slopewater region shortly downstream. However, the correlations between the upstream and downstream flows (with respect to the NESC) reach maxima of 25%, and are therefore not sufficiently high to establish a dependence of the SJ on the parameters of the incident GS (cf Fig. 22b and d).

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5. Role of the DWBC in the SJ dynamics 5.1. Effect on the mean flow The superposition of the upper and lower layer horizontal velocity fields (i.e.: weighted vertically averaged velocities above and below the pycnocline) in Fig. 23 show remarkable circulation patterns: between the seamount chain and the Grand Banks, the SJ and DWBC appear to be adjacent to each other, while flowing in opposite senses. The cores of these two currents remain adjacent for all simulated years, and averages over shorter periods. the velocity profiles of sections A and C (Fig. 2b) show that below 2000-m depth, the flow changes sign, and the DWBC extends further east under the SJ. Further offshore are the GS and the southwestward flow mentioned earlier, and identified as a portion of the Northern Recirculation Gyre (negative velocities at sections A, D between the SJ and the GS in Fig. 2b). The SJ appears, therefore, to be sandwiched between the DWBC and one of the recirculation subgyres. However, the DWBC seems to exert the dominant constraint on the SJ: note for example the SJ following the topographic bump south of the Gulf of St. Laurent channel, quite similarly to the DWBC, yet in the opposite direction. Although the DWBC follows the f/h contours along the continental slope, the SJ is a surface current with no barotropic signature near the bottom. Therefore, the tendency of the SJ core to follow the topography is more related to the DWBC, and also to the location of the recirculation cells. The MICOM slopewater velocity profile is very similar to the observations of P99 in the same area as the geostrophic velocity contours also suggest a coupling between the SJ and the DWBC: both currents are also adjacent and their boundaries overlap slightly on their common edge. The observed SJ-core has a maximum depth of 400 m and a transport of the order of 3 Sv. However, the main difference lies in the structure of the DWBC, which in the real ocean, is composed of several distinct water masses (cf Pickart, 1992). In MICOM simulations, the main DWBC component is centered at 2500 m, and corresponds to the lower North Atlantic

Fig. 23. 1980–1983 velocity fields for the ‘‘upper layer” flow, defined as the flow above the pycnocline (vertically averaged over layers 1–11) in red, and the ‘‘lower layer” flow, defined as the subpycnocline flow (layers 12–17 vertical average) in blue, superimposed on the 350- and 4000-m isobaths. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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Fig. 24. Probable ranges from EOFs of the correlation matrix for the salinity (a) and velocity (b, c) of the slopewater column downstream of the NESC. Salinity contour-labels are expressed in psu, and velocity contour-intervals are 3 cm/s. The DWBC is represented by the high salinity and negative velocity cores against the continental slope.

Deep Water (NADW), with maximum velocities extending from 2000 to 3500 m. It is topped by weaker velocities between 500 and 1500 m, which may be related to the LSW and shallow DWBC originating from the

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southern Labrador Sea (Pickart, 1992), although the 4 year mean velocities at these depths indicate a water mass flowing slightly away from the continental slope. Hydrographic surveys (P99) show that the velocity cores of the DSOW and the LSW are slightly shifted horizontally, because they flow along the continental slope at different depths. But the vertical resolution below the pycnocline in MICOM may not be high enough to differentiate between the two DWBC water masses. Previous work on the implementation of the North Atlantic water mass circulation in MICOM with lower horizontal resolution (Smith et al., 2000) pointed to the difficulty in reproducing the LSW-formation: while the Denmark Strait overflow is specified in the northern boundary region, the convective processes in the Labrador Sea responsible for the LSW-formation constitute a challenge to simulate with the present configuration of the isopycnic model. As a result, the paths of the LSW and lower NADW are not captured at distinct density levels in this simulation. Note that the streamlines composing the SJ have maximum concentration between 58°W and 50°W (Fig. 4a), indicating a transport intensification at these longitudes. The major constraint on the SJ (as displayed in Figs. 23 and 2b) appears to be the presence of the DWBC flowing southward along the continental slope on a path that is between the shelfbreak and the SJ: the strong apparent coupling of the SJ with the DWBC leads to the assumption that it may not be possible for the SJ to cross the DWBC. The fact that the two currents flow adjacent and in opposite directions results in intense vertical shear, leading therefore to a doming-up of the isopycnals. The strong tilt of the layer interfaces also acts like f/h contours, thus making it impossible for the SJ to cross these contours with relatively small layer thicknesses. Therefore in the mean, the slopewater current always appears adjacent to the DWBC, with the core slightly shifted to the east. 5.2. Variability of the SJ and DWBC downstream of the NESC An intensity coupling of the SJ with the DWBC is apparent from the 4 year-seasonal velocity averages of section A (Fig. 6), with both currents reaching maximum transports in the fall (SJ: 12.5 Sv, DWBC: 22.6 Sv) and minimum transports around winter/spring (SJ: 4.2 Sv, DWBC: 17 Sv). There is a similarity with the results of P99’s EOF-analysis, performed on the measured quantities for the slopewater column, where the minimum and maximum modal states display a coupling in intensity of the SJ and components of the DWBC, with a variation in phase of the SJ and DSOW. It remains to be seen whether this intensity coupling is direct (via vertical shear of the horizontal velocities) or indirect, such as a dependence of both currents on the GS dynamics. Nevertheless, the apparent seasonal coupling of the NE transport and DWBC between the NESC and 55°W is a remarkable feature worth a more detailed investigation with an improved model simulation. The EOFS of the salinity and projected velocity for the slopewater column in section A are calculated using the correlation matrix (cf Fig. 24) in order to monitor the interaction of the SJ with the DWBC: the dominant mode for the salinity (56% of the total variability) is actually controlled by the numerical excesses of the DWBC, due in part to the model’s salinity drift (Treguier et al., 2005), yet it shows a distinct coupling of the upper–lower layer flows: a saltier DWBC is correlated with a wider spreading of the Shelf’s fresher water under a thinner surface salinity tongue, with dominant periods of >1 year, 1 year and 8–9 months. The first two modes for the velocities also display similar behaviors including the 1 year, 9 month and 10-month periods. These modes capture in part the synchronized seasonal maximum/minimum transports in both the DWBC and SJ discussed above. 6. Summary and discussion The slopewater circulation between the NESC and the Grand Banks of Newfoundland is rich in a variety of water-masses and interactions among wind-driven and thermohaline currents. It is therefore a potentially strong element of the long-term ocean variability and its effect on climate anomaly. The main objective of this study has been to define and improve our understanding of the circulation in this region, particularly the Slopewater Jet and its connection with the Gulf Stream and the vicinity of the NESC, within the context of a realistic OGCM. Four-year outputs of a MICOM realistic high resolution simulation of the North Atlantic are used to investigate the Gulf Stream-slopewater circulation. The simulated Slopewater Jet (SJ) has characteristics similar to the observed SJ measured from hydrographic surveys between the NESC and 50°W: its path coincides with

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the sharp temperature and salinity surface fronts separating the Shelf waters from the Slope waters. The SJ water mass at the surface indicates a mixture of about 50% Shelf water and 50% Stream water, while its T, S characteristics below the mixed-layer display a GS water-mass dominance of 75–90% in each density layer, in good agreement with the observations of McLellan (1957). However, one has to take into consideration the bias introduced by the model SJ at its early stages, as it originates from occasional eddies of the GS and flows north along the continental shelf between Cape Hatteras and the NESC, whereas ocean measurements (Csanady and Hamilton, 1988) have shown that the current along the shelfbreak is actually flowing south. This discrepancy could be related to the lack of fresh water mass originating from the Labrador basin and/or river outflows in MICOM (Treguier et al., 2005). Downstream of the seamount chain, the mean SJ in the model flows adjacent to the 4000-m isobath up to the Grand Banks. The 4 year average transport of the SJ upstream of the seamount chain is 3 Sv, and it increases to 7 Sv downstream of the NESC, and then reaches a mean of 19 Sv south of the Grand Banks. While these model values between the NESC and the Grand Banks are closer to McLellan (1957) estimates than P99’s measurements of the SJ transport, the proportional transport increase is consistent with the latter’s observations between the longitudes 55°W and 50°W. Indeed, the constant rejections and eddy interactions of the GS with the SJ downstream of the NESC, contribute to this transport increase by entrainment of GS waters, and are adding to the entrainment and mixing of shelf waters. Additionally, a significant portion of the LC retroflects and mixes with the SJ, all of which end up doubling the SJ transport. Although it is overestimated by 4 Sv upstream of the NESC, it is unclear how the SJ transport is affected by this numerical bias between the NESC and the Grand Banks: the impact of the NESC on the GS should not differ significantly, the seamounts being far from the shelf, and assuming the same GS transport. One would tend to extrapolate an initial SJ originating shortly downstream of the NESC, and entraining GS waters as described in this work, with a much weaker northeast transport. However, an adequate representation of the fresh water mass means also a stronger LC transport, which should result in an increase of retroflected LC waters and shelf waters being entrained in the SJ south of the Grand Banks. The dominant mode of the SJ velocities downstream of the NESC corresponds to a net lateral translation of the current’s highest velocity core, resulting from a partial merging with the GS, and taking place predominantly with a 9-month periodicity. The 9-month period is characteristic of GS meandering intensity variability, and was related by Lee and Cornillon (1995) to the mechanism of Rossby wave interaction with the Stream. It is doubtful however, that linear Rossby waves and simple stability arguments are relevant in the highly nonlinear GS. Schmeits and Dijkstra (2000) on the other hand, find that the 9-month signal is triggered by a barotropic instability of the North Atlantic wind-driven gyres. At 55°W, P99 also find a dominant mode that captures the lateral translation of the SJ. This mode corresponds to the type of variability of the simulated SJ at section A (and mode 2 at section D, not shown). The main difference lies in the modal state of the velocity profile: the SJ derived by P99 has smaller velocities and a lateral shift that is distinct from the GS-translation (zero-velocity contour in between), although it depends on the GS, while the SJ simulated in MICOM is not fully differentiated from the GS. However, the 49% velocity eigenvalue in the model is within the same range of the density eigenvalue calculated by P99, with both capturing the GS’s strong kinetic influence on the SJ, via meandering and eddy activity. Analysis of seasonal averages shortly downstream of the NESC (section A) revealed a highest SJ transport in the fall (12.7 Sv versus 4.3 Sv in spring), consistent with the more northern path of the Stream at this time of the year, both in the model and from observations. Yet it is unclear how much of this eddy-water stays in the SJ, versus how much goes back to the GS. The major transport contribution to the SJ at this cross-section is from the GS with a dominant annual signal, whereas the transport seasonality weakens further downstream, with a more prominent 10-month signal in the model. The second velocity mode of the SJ downstream of the NESC confirms the annual transport variability related to the GS seasonality, and represented at section A by the GS’s first mode. South of the Grand Banks however, the T, S characteristics of the eastward slopewater flow still display a strong seasonality, with evidence of more LC water-mass entrainment and mixing in the spring and summer (due to the LC retroflection taking place in winter/spring). The first mode at 50°W of the upper-slopewater explains 37% of the total variability and captures the transport variations of the SJ and LC, with predominant timescales of 10 months, 1 year and >1 year. EOF-derivations of each current separately, affiliate these

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dominant periods with the intrinsic frequencies of the LC (1 year), the SJ (10 months, >1 year) and the GS (>1 year) in the numerical model, which infers that the variability of the SJ transport at 50°W results primarily from the interaction of GS-meanders with the upstream SJ. Observations of P99 also yield an upper slopewater non-dominant mode (similar eigenvalue) south of the Grand Banks, capturing the same type of variability. P99 assumed that the presence of the LC which adds another major source of variability in this region, explains the lack of dominant mode for the SJ, and this is confirmed by the different intrinsic frequencies of the SJ and LC derived in this study. The presence of the NESC appears to contribute to the advection of GS waters inshore, and subsequent increase in transport and persistence of the SJ in MICOM. Analyzes of the mean flow and kinetic energy of the whole water column revealed the existence of Taylor caps, which appear to be the main mechanism captured by the numerical model to account for the strong upper ocean-topographic coupling. The spatial distribution of the seamount chain, the proximity of the continental slope, and the jet-related distortion of the pycnocline redistribute the potential vorticity of the deep flow via the lower layer isopachs, in the form of recirculation cells, which in turn affect the upper-layer circulation by triggering a fan-shape streamline divergence of the GS. While a northern portion of the GS is deflected into the slopewater region and contributes to the SJ transport along the 4000-m isobath, the main portion of the Stream is constrained to meander anticyclonically around the Gregg–San Pablo subchain. The streamline concentrations occurring in these two distinct directions point to a Gulf Stream’s spatial bifurcation, rather than a mere reduction of the jet’s coherence. The major change of GS/SJ dynamics occurs at the NESC, as highlighted by the variability of the pycnocline’s depth: the first EOF of the upper-layer thickness above the NESC displays a qualitative dependence of the fan-shape streamline dispersion on the strength and intersecting latitude of the incident GS. As the 9-month and 5–6-month mesoscale signals of this subspace get amplified in the area above the seamount chain, it appears that the NESC bathymetry could enhance the GS-meandering activity in MICOM. Although the anticyclonic meander of the GS around Gregg–San Pablo has been observed in the ocean, there are no available measurements in the vicinity of the NESC to support the GS-bifurcation scenario. It has been suggested that OGCMs tend to overestimate the NESC effect on the upper-layer flow (Ezer, 1994; Blayo et al., 1994). The mean path of the SJ from the NESC up to the Grand Banks is remarkably adjacent to the path of the DWBC. This geographic coupling (also present in P99’s observations at 55°W and 50°W) seems related to the strong vertical shear of the two currents flowing in opposite directions, leading to a tilting of the isopycnals and preventing the SJ path from shifting further west. The phase in transport variation is captured by the full slopewater EOF of the velocity, temperature, and salinity in MICOM. Although P99 find a similar coupling of the SJ with the DSOW, the transport seasonality of the DWBC in the model is not supported by observations, and the salinity drift in MICOM related to the difficulties in producing and maintaining the correct watermass production in recent OGCMs (Treguier et al., 2005) affects significantly the statistical modes. Further understanding of the GS-slopewater system will require a more systematic observation of the slopewater circulation, in order to differentiate the interannual from the subannual/seasonal signals, and monitor the GS behavior above the NESC. Numerical models on the other hand, have to overcome two difficult challenges to address the issues raised in this study: first, a very accurate simulation and/or parameterization of the deep water-mass formation processes involved in the thermohaline circulation, and second: an adequate representation of the full dynamics (mixing and internal waves) generated by the interaction of the flow with seamounts. Acknowledgements This work was supported by the Office of Naval Research Grant N00014-03-1-0284. The authors thank ¨ zgo¨kmen, Zulema Garraffo, Bill Johns, and two anonymous reviewers for their helpful comments. Tamay O References Adamec, D., 1988. simulation of the effects of seamounts and vertical resolution on strong ocean flows. J. Phys. Oceanogr. 18, 258–269. Blayo, E., Verron, J., Molines, J.M., 1994. Assimilation of TOPEX/POSEIDON altimeter data into a circulation model of the North Atlantic. J. Geophys. Res. 99, 24691–24705.

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