Diagnosing and Picturing the North Atlantic Segment ... - AMS Journals

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JOURNAL OF PHYSICAL OCEANOGRAPHY

VOLUME 32

Diagnosing and Picturing the North Atlantic Segment of the Global Conveyor Belt by Means of an Ocean General Circulation Model BRUNO BLANKE, MICHEL ARHAN, SABRINA SPEICH,

AND

KARINE PAILLER

Laboratoire de Physique des Oceáns, CNRS-IFREMER-UBO, Brest, France (Manuscript received 13 February 2001, in final form 25 September 2001) ABSTRACT The monthly mean velocity, salinity, and temperature fields of a numerical simulation of the World Ocean climatological circulation are used to study the intensity and pathways associated with the meridional overturning in the North Atlantic. Lagrangian diagnostics based on the computation of several hundreds of thousands of individual three-dimensional trajectories are combined with an appropriate study of water mass potential densities in order to describe the warm and cold limbs of the so-called conveyor belt. Circulation schemes are established for both limbs of the overturning, and can be easily compared with schemes or transport estimates deduced from direct measurements, as the model temperature and salinity fields are constrained to remain close to the observed climatology. Diagnostics emphasize most typical pathways as well as main mass transfers that lead to the establishment of such numerical circulation schemes.

1. Introduction North Atlantic Deep Water (NADW) is one of the major deep-water masses in the World Ocean. Formed in the North Atlantic at high latitudes, it plays a crucial role in the meridional heat budget as it flows southward at deep immersions (1000–4000 m), being necessarily balanced in volume by an equivalent northward flow of surface (0–1000 m) warm water. The intensity, distribution, and variability of this warm limb of the global conveyor belt (Gordon 1986; Broecker 1991; Speich et al. 2001) are known to influence the European and, more generally, the global climate (Rahmstorf 1997; Stocker and Schmittner 1997). The general denomination NADW represents a complex aggregate of several distinct dense water masses, formed or converted in different locations. Classically, three origins are distinguished (Smethie 1993; Dickson and Brown 1994; Fine 1995). Labrador Sea Water (LSW), the least dense of the three components, is formed by deep wintertime convection in the central Labrador Sea. Denmark Strait overflow water (DSOW), the densest component of NADW, enters the western North Atlantic through Denmark Strait, between Greenland and Iceland, and locally entrains subpolar water (by overflow-enhanced mixing). The third classical component, Iceland–Scotland overflow water (hereafter ISOW), enters the North Atlantic mainly through the Corresponding author address: Dr. Bruno Blanke, Laboratoire de Physique des Oceáns, UFR Sciences et Techniques, 6 avenue Le Gorgeu, B.P. 809, 29285 Brest Cedex, France. E-mail: [email protected]

q 2002 American Meteorological Society

Faroe Bank Channel, between Iceland and Scotland. It locally entrains ambient subpolar water and eventually flows through the Charlie-Gibbs Fracture Zone (hence its other designation, GFZW) around the Reykjanes Ridge toward the western part of the basin, where it is sandwiched between LSW and DSOW. Pickart (1992) was the first to evoke a possible fourth source: a water mass formed by convection in the southern Labrador Sea, but less dense than LSW, and called Shallow LSW or Upper LSW (Molinari et al. 1992; Rhein 1994; Smethie 1993; Smethie et al. 2000). The origins of the warm limb of the global conveyor belt are the topics of lasting debates among oceanographers. According to a warm route hypothesis (Gordon 1986), the Pacific and Indian oceans are linked to the upper Atlantic with an exchange of warm water south of Africa. In a cold route conjecture (Rintoul 1991), the dominant contribution of water and heat into the Atlantic is obtained directly at Drake Passage, south of America. A third likely route was also derived recently from model results, linking the tropical Pacific to the tropical Atlantic through a path south of Australia (Speich et al. 2001). The fate of these waters within the North Atlantic is less controversial since they must balance in volume the production of deep water occurring in the northern Atlantic or within the Mediterranean Sea. A precise diagnostic of the mass transfers and associated pathways can, however, prove difficult to obtain, considering the nonstationarity of the deep-water formation process, either on seasonal or interannual timescales (e.g., Dickson et al. 1990; Curry et al. 1998). The description of the circulation in the North At-

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lantic that we currently have relies mostly on extensive collections of accurate direct or indirect measurements. Various authors have been successful in producing circulation schemes, merging all possible sources of data and trying to match any available piece of knowledge (e.g., Schmitz and McCartney 1993; Macdonald and Wunsch 1996; Schmitz 1996; Sloyan and Rintoul 2000; Smethie et al. 2000). We focus here on a more general problem that arises in physical oceanography: Where does the water come from and where does it go? Which conversions happen along its path? Pathways and flows (usually expressed in Sverdrups: Sv [ 10 6 m 3 s 21 ) are the essence of such circulation schemes, and are keys to a full understanding of the global ocean circulation. Historically, however, ocean models have dealt with the ocean circulation in terms of Eulerian quantities: velocity components, temperature, salinity, or any other tracers are most often computed on discrete points of a given three-dimensional (3D) mesh, for successive time steps. Most concepts related to quantitative Lagrangian diagnostics applied to numerical ocean models have therefore remained rather vague until the development of intensive computations of individual trajectories aiming at quantifying interbasin connections or water mass displacements (Doö¨s 1995; Blanke and Raynaud 1997; Blanke et al. 1999, 2001; Speich et al. 2001). They are now being extensively developed in the context of an European Economic Community–funded MASTIII project, TRACMASS, whose main goal is tracing the water masses formed in the North Atlantic and the Mediterranean Sea. As a contribution to TRACMASS, we show in this paper that the output of a complex ocean general circulation model (OGCM) can be analyzed with Lagrangian techniques to produce circulation schemes equivalent to those inferred from direct measurements. These diagnostics may help or complete the physical understanding of the 3D NADW movement and associated warm water transfer within the North Atlantic. This study will thus investigate the modeled southward spreading of the dense waters formed by winter convection in the subpolar gyre and in the Arctic Seas or linked to the Mediterranean outflow. Most circulation schemes usually emphasize the western boundary structure of the deep southward flow (Schmitz and McCartney 1993) whereas several studies show the presence of NADW in the eastern Atlantic (Talley and McCartney 1982; Paillet et al. 1998). We will estimate the relative contributions of the major Arctic straits and of the Gulf of Cadiz to the North Atlantic overturning. The way eastern Atlantic NADW merges with waters formed in the Mediterranean Sea or entrained vertically at the level of the Mediterranean outflow will be addressed. The next section of the paper details the modeling tools used in the study (OGCM and Lagrangian diagnostics). Section 3 documents the intensity of the meridional overturning related to the existence of the conveyor belt in the North Atlantic. We first calculate the annual mean streamfunction of zonally integrated vol-

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ume transport. Then we run an initial Lagrangian experiment (with hundreds of thousands of particles) documenting the whole northward inflow at 5.38N and we isolate the particles that correspond to a genuine conversion of warm into cold water within the North Atlantic. Sections 4 and 5 detail both in a quantitative and qualitative way the pathways obtained in the North Atlantic for the warm limb and the cold limb of the Atlantic conveyor belt, respectively. The Lagrangian analysis focuses on specific sectors of the domain, discriminating the subpolar and subtropical segments of the overturning. A discussion of the horizontal distribution of the overturning is given in section 6 and sources of error for our Lagrangian diagnostics are studied in section 7, before a conclusion is drawn in a last section. 2. Modeling tools a. Ocean model The OGCM used to diagnose the origins and pathways of ocean water masses is the Oceán Paralle´lise´ (OPA) model (Madec et al. 1998) in its global configuration. The domain extends from the Southern Ocean at 788S to 908N. The singularity of the North Pole is removed by introducing an appropriate coordinate transformation that includes a double, numerical inland pole. As a result, the zonal resolution is 28 within the whole southern hemisphere and only slightly distorted in the Northern Hemisphere. The meridional grid interval varies from 0.58 at the equator, to a maximum of 1.98 in the Tropics. There are 31 levels in the vertical, with the highest resolution (10 m) in the upper 150 m. The bottom topography and the coastlines are derived from a global atlas, based on Smith and Sandwell’s (1997) study completed by values from the 59 3 59 Earth Topography 5-Minute Gridded Elevation Dataset (ETOPO5) for the northernmost latitudes of the domain. A zoom over the North Atlantic of the exact global topography of the model, without any interpolation or smoothing, is shown in Fig. 1 together with the sections that will be used for our Lagrangian quantitative experiments. The model is forced by a daily climatology obtained from the European Centre for Medium-Range Weather Forecasts (ECMWF) 15-yr (1979–93) reanalyses, and smoothed by an 11-day running mean. The experiment was designed to recover and study the dynamics associated with the observed global ocean hydrography. Therefore, our analyses are carried out over the last 12 months of a ten-year simulation in which a restoring term to the Levitus (1982) climatology was added to the potential temperature and salinity equations (Madec and Imbard 1996). This Newtonian damping acts everywhere except in the 208S–208N latitude band and in the surface mixed layer (not to interfere with model’s fast adjustment in these regions). Its intensity is defined as the inverse of a timescale that varies from 50 days in the uppermost ocean layers to 360 days down

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FIG. 1. Model topography for the North Atlantic. The sections used in the various Lagrangian experiments (see Table 1) are indicated with ‘‘star’’ symbols. The gray palette is a linear function of depth. Annual mean mass exchanges across mentioned sections are given in Sv. The plot uses a vertical, satellite view projection.

to 5000 m, to account for slower dynamic adjustments. The constraint is also slightly relaxed poleward of 508 N and S because of sparser observations. The damping comes progressively to zero within 1000 km of the coastlines as boundary currents may not well be captured by the climatologies. The model physics [isopycnal lateral mixing, 1.5-order vertical turbulent closure scheme (Blanke and Delecluse 1993), Beckmann and Do¨scher’s (1997) parameterization for overflows] is able to recover the boundary and equatorial currents not well resolved in the Levitus (1982) climatology. The relationship used in the model to compute density from potential temperature and salinity is the equation of state proposed by Jackett and McDougall (1995). b. Lagrangian analysis The equations of the OGCM are discretized on a C grid. Finite differencing on this grid expresses mass conservation in a very natural way as it states that entering mass in a given gridcell must be balanced by an exactly equal outflow. This description is ideal for the computation of analytical mass-preserving streamlines (Blanke and Raynaud 1997) for a given sampled velocity field. Assuming the velocity field to be constant over periods equal to the available sampling time of the OGCM output, one may associate successive segments of such streamlines to actual 3D trajectories of fictive fluid particles being advected by the general circulation computed by the OGCM. The notion of a mass-pre-

serving scheme for computing trajectories is essential when using it for water mass tracing, since it will ensure that the calculated trajectories will be able to mimic the genuine movement of water mass parcels, without intercepting the coastline or ocean bottom. The next step in our Lagrangian analysis is to use this trajectory scheme to advect particles describing a whole (or fraction of ) water mass. As proposed by Doö¨s (1995) and adapted by Blanke and Raynaud (1997) this is achieved by inseminating a water mass on a given geographic section with many (hundreds of thousands of ) particles, each of which is associated with an infinitesimal fraction of the transport of the selected water mass. Infinitesimal volumes of water will be conserved along model streamlines and for selected final destinations (another geographic section, or the fulfillment of a hydrographical criterion) infinitesimal transports may be added, and directional transports can be produced. Online diagnostics could be run within the OGCM, but it is more efficient to develop offline Lagrangian calculations, based on successive outputs of the model, like monthly or daily outputs, depending on the complexity in time of the flow to analyze. Offline diagnostics allow backward computations of trajectories (simply by multiplying all velocity outputs by 21, and reversing their order). The joint use of backward and forward Lagrangian calculations also permits estimation of the error made in computing directional transports (see section 7). Any transport (from section A to section

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TABLE 1. Detailed characteristics of all Lagrangian experiments (direction of integration, initial conditions, initial and final interception sections). The origin of the set of particles in use is also specified when obtained from a former Lagrangian experiment. Initial sections also define final sections (by intercepting recirculating particles). Sections used for interception Expt

Integration

1

Forward

2a

Forward

2b

Forward

3

Forward

4a

Backward

4b

Backward

5

Forward

Particles in use

5.38N

5.38N Whole northward flow From expt 1 Selected initial positions at 5.38N From expt 2a Final positions at 538N Bering Whole northward flow From expt 1 Selected final positions at 5.38N From expt 4a Final positions at 538N From expt 1 Selected initial positions at 5.38N

Initial

B) can indeed be calculated in two different and independent ways: inseminating A and summing the transports of the particles that do reach B, or inseminating B and summing the transports of the particles that do originate in A. Moreover, offline Lagrangian diagnostics permit to loop over a climatological year while calculating trajectories without the constraint of the true length of the OGCM simulation for setting up the limits of the Lagrangian integration. Finally, circulation schemes are obtained by computing the horizontal streamfunction related to the vertical integration of the 3D transport field determined by the displacement of the particles used to describe the flow of any selected water mass (Blanke et al. 1999;

Initial

98W

538N

ISOW

DSOW

3 3

3 Initial

3

3

3 Initial

Bering

Initial 3

3 Initial

3

3

Initial

Speich et al. 2001; Blanke et al. 2001). The full characteristics of the various Lagrangian experiments analyzed in the following sections are detailed in Table 1. 3. Intensity of the North Atlantic meridional overturning a. Overturning streamfunction We first diagnose the overturning streamfunction in the North Atlantic (including the Mediterranean Sea) by integrating zonally the meridional and vertical velocity fields and averaging the result over the whole annual cycle (Fig. 2). The amplitude of the overturning in the

FIG. 2. Annual mean zonally integrated circulation. The contour interval is 1 Sv. Dotted negative values refer to an anticlockwise circulation. An expanded scale is used in the first 500 m.

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TABLE 2. Water mass definitions at 5.38N used for the study (see Lux et al. 2001). Indices for s represent the reference level (in 10 3 db) for the calculation of potential density. Water mass name

Abbreviation

Lower limit

Upper limit

Warm limb Surface Water Thermocline Water Lower Thermocline Water Intermediate Water

SW TW LTW IW

Upper Deep Water Middle Deep Water Lower Deep Water

UDW MDW LDW

Bottom Water

BW

s0 s0 s0 s1

5 5 5 5

24.58 26.75 27.15 32.20

Ocean surface s 0 5 24.58 s 0 5 26.75 s 0 5 27.15

Cold limb

North Atlantic diagnosed in z coordinates is 16.9 Sv at 2.58N, 16.6 Sv at 208N, and 16.4 Sv at 408N centered around 1200 m. The net meridional flow throughout the Atlantic is southward and amounts to 0.8 Sv as the model allows the connection of the Atlantic and the Pacific through Bering Strait. Apparent locations of strong downward movements include the center of the subtropical gyre (around 308N), the subpolar gyre (from 508 to 608N) and higher latitudes in the Greenland and Norwegian Seas. The pinching of the overturning cell at 368N is linked to water mass modification within the Mediterranean Sea (strong convection) and the Gulf of Cadiz (entrainment of surface and subsurface water to intermediate levels). The counterrotating cell diagnosed deeper than 3500 m is the signature of the northward transport of bottom water north of the equator (Bo¨ning and Herrmann 1994). b. Water mass conversions The zonally integrated description of the circulation given by the meridional overturning streamfunction cannot depict the full complexity of horizontal water mass movements in the North Atlantic. We will use convenient Lagrangian diagnostics to complement this vertical view by defining and tracing in the model the cold and warm limbs of the North Atlantic segment of the global conveyor belt (NACB hereafter). We document the whole model northward flow at 5.38N with several hundreds of thousands of particles, each of which is associated with a small fraction of the incoming transport (expt 1, see Table 1). We compute the full set of trajectories within the North Atlantic and Arctic Oceans. As the Bering Strait in the model only lets water flow from the Pacific toward the Arctic Ocean, the unique possible fate for all particles is a return journey back to 5.38N. We adopt the exact same definitions for water masses at 5.38N as those used by Lux et al. (2001) at 7.58N when studying the interhemispheric exchanges of mass and heat in the Atlantic Ocean. Their limits in potential densities are given in Table 2. With notations identical to those proposed by Lux et al. (2001), we distinguish seven water masses: Surface Water (SW),

s 2 5 36.98 s 4 5 45.83 s 4 5 45.90

s1 5 32.20 s 2 5 36.98 s 4 5 45.83

Ocean bottom

s 4 5 45.90

Thermocline Water (TW), Lower Thermocline Water (LTW), Intermediate Water (IW), Upper, Middle, and Lower Deep Water (UDW, MDW, LDW, respectively) and Bottom Water (BW). All our diagnostics will document and name water masses according to their definition given at 5.38N, regardless of the subsequent conversions undergone along the trajectories integrated forward or backward in time. For our study and consistently with Lux et al. (2001), the cold limb of the NACB will be defined by the particles flowing southward in the tropical Atlantic (at 5.38N) with a potential density comprised between 45.90 (s 4 referred to 4000 db) and 32.20 (s1 referred to 1000 db) and originating initially in a northward warm flow at the same latitude with a potential density less than s1 5 32.200, namely the warm limb of the same NACB. This Lagrangian definition lies on a careful analysis of the hydrography related to each individual trajectory calculation. This allows us to focus both on the cold and warm limbs of the overturning as described by the very same set of particles, each of which is related to true NADW formation north of 5.38N and therefore unrelated to meanders within the domain or conversions within the NADW layer. Table 3 sums up all the transfers diagnosed from 5.38N back to 5.38N, classifying the particles (and the mass transport they explain) according to their initial and final hydrographical state (based on model potential density). Most of the total northward inflow consists of SW (31% of 52.8 Sv). Other dominant contributions come from IW (15%) and LTW (12%), and even UDW (13%) and MDW (11%), as the initial section intersects some recirculation of these water masses in the near equatorial Atlantic. Similarly, most of the outflow is made of UDW (29%), MDW (24%), and LDW (12%), but also of SW (16%) as the near equatorial section intercepts the surface equatorial and subequatorial gyres. Restricting our analysis to the warm and cold limbs of the NACB modifies this distribution as water masses that do not participate truly in the conversion of warm waters into NADW are no longer included in the diagnostic. Half of the warm limb transport consists of SW (49% of 16.8 Sv), whereas TW, LTW, and IW contribute to 9%, 20%,

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TABLE 3. Water mass conversions within the North Atlantic (excluding Bering Strait waters, 0.8 Sv). All mass transports are given in Sv and are obtained by summing the individual transports of the particles from expt 1 achieving the same kind of transfer. Bold figures are related to our definition of the cold and warm limbs of the NACB, i.e., the conversion of water lighter than s1 5 32.20 into water denser than s1 5 32.20. Final state (5.38N) SW

TW

LTW

IW

UDW

MDW

LDW

BW

Total

Warm limb

Initial state (5.38N) SW 7.9 TW 0.6 LTW 0.1 IW 0.0

Water

0.3 2.7 0.1 0.0

0.0 0.1 2.2 0.0

0.2 0.0 0.2 4.0

4.8 0.9 1.7 2.4

2.7 0.5 1.3 1.1

0.7 0.1 0.4 0.2

0.0 0.0 0.0 0.0

16.6 4.9 6.0 7.7

8.2 1.5 3.4 3.7

0.0 0.0 0.0 0.0 3.1

0.0 0.0 0.0 0.0 2.3

0.0 0.0 0.0 0.0 4.4

5.4 0.1 0.0 0.0 15.3 9.8

1.1 5.5 0.2 0.2 12.6 5.6

0.1 0.4 4.0 0.6 6.5 1.4

0.0 0.0 0.0 0.0 0.0

6.6 6.0 4.2 0.8 52.8

UDW MDW LDW BW Total Cold limb

0.0 0.0 0.0 0.0 8.6

and 22%, respectively. This distribution is very similar to the net transports diagnosed at 7.58N by Lux et al. (2001), namely 53%, 7%, 21%, and 19% for SW, TW, LTW, and IW, respectively, for a net 21.9 Sv overturning. Our water mass breakdown for the cold limb consists mostly of UDW (58%) plus some MDW (34%) and a small amount of LDW (8%). It differs a lot from the percentages proposed by Lux et al. (2001) for the net southward NADW export (15%, 15%, and 70%, respectively) or by Friedrichs and Hall (1993), who diagnosed at 118N a net southward NADW flow of only 7.4 Sv, including a net northward MDW contribution. The intensity of the overturning, as defined in Table 3 from true water mass conversions (16.7 Sv), results in close agreement with the value diagnosed from the meridional mass streamfunction. It matches the apparent water mass consumption or formation, which we can diagnose from purely Eulerian diagnostics applied at 5.38N calculating the difference between entering and exiting individual water masses (Table 4). Total northward or southward transport contributions for individual water masses may be underestimated because of a poor account for small-scale movements and recirculations. TABLE 4. Local (5.38N) NACB warm and cold limbs estimates. Inflows and outflows at 5.38N are expressed in Sv and are derived from Table 3. The net budget for a water mass at 5.38N is an estimate of its transformation within the North Atlantic and can be used to evaluate the intensity of the NACB warm and cold limbs, respectively. Northward inflow Water mass at 5.38N SW TW LTW IW UDW MDW LDW BW

16.6 4.9 6.0 7.7 6.6 6.0 4.2 0.8

Southward outflow at 5.38N

Net budget at 5.38N

8.6 3.1 2.3 4.4 15.3 12.6 6.5 0.0

8.0 1.8 3.7 3.3 28.7 26.6 22.3 0.8

Warm and cold limbs estimates 16.8

17.6

16.8

The model indeed suffers from a lack of resolution in the lateral, vertical, and time directions. Such estimates may not compare well for instance with direct measurements of the tropical Atlantic circulation. The Lagrangian definition adopted in Table 2 provides more accurate results since the implied conversions necessarily involve a genuine large-scale displacement in the Atlantic domain, and not only very local recirculations that the model may be unable to account for. Therefore, we will use this definition to develop subsequent Lagrangian diagnostics based on a proper discrimination of water masses and their conversions in the North Atlantic and Arctic Oceans. The net flow we diagnose in the model across 5.38N in the NADW layer is southward and amounts to 17.6 Sv (Table 4). Most of the NADW circulation is concentrated near the western boundary, and the meridional transport integrated eastward peaks at 22 Sv (southward flow) at 408W. This value is in the medium range of in situ estimates from direct velocity or mass field measurements as, for instance, summed up for the near equator and at the western boundary by Rhein et al. (1995), with values ranging from 17 to 25.8 Sv within the 58S– 108N band. The combined influence of several parameters explains the variation in transport computations, as the geographical location [latitude, zonal extent adopted for the definition of the deep western boundary current (DWBC), the methodology (use of geostrophic velocities or direct velocity measurements, choice for reference levels), and the period of observation (plausible seasonal and interannual variability of the NADW flow). One must note here that our definition of cold limb water includes the NADW formed in the Arctic Sea as well as formation of deep water in the Mediterranean basin, or entrainment of central or intermediate waters within the subtropical Atlantic. This latter component eventually joins the DWBC in the southern part of the subtropical gyre. Our estimate compares well with the evaluations found in the literature for the deep-water

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FIG. 3. Horizontal mass streamfunction related to the vertically integrated transport of the NACB [see Blanke et al. (1999) for details about the calculation] as computed in expt 1. The value of the streamfunction is arbitrarily set to 0 above Africa. Negative values are dotted. The contour interval is 2 Sv.

formation in the North Atlantic. McCartney and Talley (1984) indeed propose 14 Sv as the rate of deep water formation by convection in the northern North Atlantic. Dickson and Brown (1994) evaluate a NADW transport of 13.3 Sv south of Cape Farewell, a quantity they consider consistent with recent estimates of the net abyssal southward flow through the North and South Atlantic. At the lower end, Friedrichs and Hall (1993) suggest 12 Sv as the intensity of the meridional overturning cell. 4. Picturing and quantifying the warm limb of the NACB Annual mean mass fluxes across specific sections of the North Atlantic domain are given on Fig. 1. The exchanges diagnosed at the level of the Arctic straits are in fair agreement with published values as much for the warm inflow to the Nordic Seas (Dickson and Brown 1994; Schmitz 1996) as for the export of dense waters to the Atlantic Ocean (Krauss 1995). The flow across 98W is balanced since the Mediterranean Sea is a closed basin and mass is conserved in the model. We could not compare the model total eastward transport at 98W with any climatology. However we found a fair agreement between the total eastward transport at 138W integrated in the model between 328 and 398N and from the surface to s1 5 32.20 (roughly 500 m), and the value inferred by Arhan et al. (1994) for the experimental ‘‘Bord-Est’’ program (6.1 and 7.6 Sv, respectively). This eastward transport feeds the inflow into the

Mediterranean Sea, the downward entrainment, the coastal Ekman drift, and the Canary Current. Figure 3 displays the horizontal mass transport streamfunction built from the ensemble of particles that are part of the 16.8 Sv conversion of warm water into NADW. The calculation of this streamfunction is detailed in the appendices of former studies (Blanke et al. 1999, 2001): the passage of each particle, with its associated transport, is recorded and algebraically summed on each velocity point of the 3D grid. The subsequent 3D nondivergent transport field can be integrated in the vertical and pictured as a horizontal streamfunction. Figure 3 looks very similar to the standard streamfunction we could diagnose from the barotropic flow (not shown) except for the Atlantic pathways related to Bering Strait waters. Anticyclonic circulation in the subtropical gyre and cyclonic circulation in the subpolar and subequatorial regions are the dominant patterns of the flow. The complex tropical circulation is also apparent, with the retroflection of the North Brazil Current and the linked cyclonic and very narrow northern tropical cell including the North Equatorial Countercurrent (NECC). However, the vertical projection wipes out the meridional (or zonal) exchanges that occur at a same place but at different levels. Thus the figure cannot depict the full intensity of the intergyre transfers, or the ventilation of the Mediterranean basin, except for a small cyclonic patch located in the Gulf of Cadiz, inherent to the distinct directions for inflow and

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TABLE 5. Diagnosed fates for the NACB warm limb initially considered at 5.38N (expts 2a and 2b). All mass transports are expressed in Sv. Expt 2a particles intercepted at 538N (namely 9.9 Sv) are used as initial particles for expt 2b. Epxt 2a Expt 2b Recircu- Gulf of Total at lation Cadiz 5.38N to 5.38N (98W) SW TW LTW IW Total

8.2 1.5 3.4 3.7 16.8

1.1 0.1 0.2 2.0 3.4

2.4 0.5 0.4 0.2 3.5

Iceland– Recircu- Scot- Denlation land mark 538N to 538N passage Strait 4.7 0.9 2.8 1.5 9.9

3.0 0.4 1.0 0.7 5.1

1.3 0.4 1.6 0.7 4.0

0.4 0.1 0.2 0.1 0.8

outflow at Gibraltar. More selective Lagrangian experiments, using the same ensemble of particles but distinct intercepting sections, will allow us to investigate more thoroughly the pathways associated with the warm limb of the NACB. a. Warm limb circulation south of 538N The next step in our Lagrangian analysis is a separate examination of the warm and cold water flows in the North Atlantic, based on the hydrographical properties of the particles diagnosed on their initial and final positions at 5.38N. Retaining the exact same set of initial particles as in experiment 1, trajectories are now integrated forward in time until they reach 538N (a latitude close to the median position of the subpolar gyre), 98W between Spain and Africa, or until they go back to 5.38N (see Fig. 1). We will hereafter refer to this configuration as experiment 2a (see Table 1). The calculations are made individually for each component of the warm inflow (SW, TW, LTW, and IW), and the pathways from 5.38 to 538N or 98W are displayed by means of a horizontal mass streamfunction. The calculation is no longer related to the displacements of the full set of particles, but is now restricted to those corresponding to a selected initial water mass at 5.38N, transmitted to a selected destination (538N or 98W). The interception at 98W allows an easy representation of the warm water flow that may eventually reach the Mediterranean basin. Because of possible vertical mixing with Gibraltar overflow waters and because of an overall crude model representation of the topographic or coastal constraints on the strait inflow and outflows, we chose to intercept the particles at the exit of the Gulf of Cadiz, and not further east. The mass transfers we calculate are summed up in Table 5. The SW flow to 538N (4.7 Sv, Fig. 4a) exhibits some intense recirculation within the subtropical gyre (almost three times the transmitted flow). This recirculation involves a pathway north of the West Indies but also some flow within the Caribbean Sea, the Gulf of Mexico and

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then through the Straits of Florida. The pathways in the immediate neighborhood of the initial section display some eastward export of SW, compatible with the seasonal retroflection of the NECC [see Bourles et al. (1999) for direct measurements or Blanke et al. (1999) for model results]. The North Atlantic Current (NAC) brings SW to 538N in the eastern Atlantic (east of 258W). A separate examination of the water transmitted to the Gulf of Cadiz (2.4 Sv) indicates less relative recirculation in the gyre and a larger relative eastward export of water within the NECC (not shown). The pathways in the northeast Atlantic display a strong southward or southeastward curvature west of the Iberian Peninsula (including the Portugal Current, see Krauss 1986), before joining the Gulf of Cadiz south of 358N, as a cornering between the northeastward flowing NAC and the eastward flowing Azores Current (Sy 1988; Paillet et al. 1998). Finally 1.1 Sv of SW return to 5.38N without having reached any other limit. The spatial characteristics of the TW flow (Fig. 4b) are very similar to those of SW. The intensity is however much reduced (0.9 Sv to 538N and 0.5 Sv to the Gulf of Cadiz), as the warm limb at 5.38N does not involve much TW: most of the Southern Hemisphere waters have been upwelled and converted into SW in the equatorial domain, as diagnosed for instance by Blanke et al. (1999). The LTW flow (Fig. 4c) is related to weak recirculation in the subtropical gyre (as it is converted progressively to deeper water masses, and results no longer sensitive to the direct action of the surface wind stress), to a 2.8 Sv transfer to 538N and some eastward transmission to the Gulf of Cadiz (0.4 Sv), in a way similar to TW. Thermocline water and, to a lesser extent, LTW enter the domain at 5.38 in form of two separate flows: a western boundary current (the North Brazil Current) and, in the eastern tropical Atlantic, an extension and a poleward turning of the North Equatorial Undercurrent. As TW, LTW does not recirculate much back to 5.38N. The pathways for IW (Fig. 4d) do not exhibit any zonal extension in the equatorial domain. The flow is divided equally between a Gulf of Mexico route and a more direct passage along the northern coast of the West Indies. No recirculation appears in the subtropical gyre. A large fraction of the flow (2.0 Sv) recirculates back to 5.38N before reaching 538N (1.5 Sv) or 98W (0.2 Sv). As already specified, our definition for water masses is based on characteristics specified at 5.38N. During transmission to 538N or 98W, or recirculation to 5.38N, the physical properties of the flow (temperature and salinity) are modified through turbulent mixing. Hence the waters reaching 538N may have a salinity and a temperature that differ from their initial value at 5.38N. Table 6 documents these transformations by diagnosing the mean temperature and salinity and associated standard deviations for SW, TW, LTW, and IW at their initial location (5.38N) and on the intercepting sections (538N, 98W and 5.38N). We observe, in particular, that the


FIG. 4. Same as Fig. 3 but for the warm waters originating in 5.38N and transmitted to 538N or to the Gulf of Cadiz (98W) as computed in expt 2a: (a) SW with a 0.5 Sv contour interval, (b) TW with a 0.25 Sv contour interval, (c) LTW with a 0.5 Sv contour interval, and (d) IW with a 0.25 Sv contour interval.

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TABLE 6. Mean temperature and salinity properties for SW, TW, LTW, and IW participating in the NACB as diagnosed at 5.38N (where the water masses are defined) and after transmission to 538N, 98W, or 5.38N. The mean statistics (average and standard deviation) are calculated from the ensembles of particles that describe each water mass (expt 2a), by interpolation of the model temperature and salinity fields at each particle position. Temperature and salinity values are given in 8C and psu units, respectively. 5.38N

Recirculation to 5.38N

Gulf of Cadiz (98W)

538N

SW u S

26.5 6 0.6 35.47 6 0.79

3.6 6 0.5 34.98 6 0.03

14.1 6 1.2 36.00 6 0.16

8.8 6 1.2 35.27 6 0.09

TW u S

14.8 6 3.4 35.43 6 0.36

3.6 6 0.5 34.99 6 0.03

14.2 6 11 36.01 6 0.15

9.2 6 1.3 35.29 6 0.09

LTW u S

8.8 6 1.3 34.78 6 0.11

3.5 6 0.6 34.98 6 0.03

14.5 6 1.0 36.05 6 0.15

9.3 6 1.2 35.28 6 0.08

IW u S

5.4 6 0.7 34.65 6 0.07

3.7 6 0.6 34.96 6 0.04

15.6 6 1.0 36.21 6 0.13

8.7 6 1.7 35.23 6 0.10

northward IW flow in the North Atlantic is either converted into lighter water masses before it reaches 538N or entrained to deeper (NADW) levels and recirculate back to 5.38N. The presence of a northeastward path of IW in the model (Fig. 4d) matches the finding of Tsuchiya (1989) that a part of the IW is conveyed to high latitudes by the Gulf Stream and the North Atlantic Current. Tsuchiya rested his conclusion on the presence of a high silicate tongue approximately coincident with the path shown in Fig. 4d. At variance with the northward extension of the silicate signal, he noted the disappearance of the IW salinity signature at about 208N, which is also compatible with the property changes reported in Table 6. A circulation schematics in Schmitz and McCartney (1993, their Fig. 14b) illustrates the two possible fates of the IW entering the North Atlantic, either recirculation to the south of 258N or further northward continuation (2 Sv in their cartoon) in the Gulf Stream/North Atlantic Current system. In the model, the water transmitted to 538N that is defined as IW at 5.38N and that participates in the NACB amounts to 1.5 Sv (Table 5). It is warmer at 538N (8.78C on average) than at 5.38N (5.48C on average), being progressively mixed with overlying waters (see Table 6). The fraction of the NACB that reaches 538N no longer supports the water mass decomposition adopted at 5.38N: model temperature and salinity, interpolated on the particle’s final positions at 538N, look very similar for all water masses defined initially at 5.38N. The fraction of IW that recirculates back to 5.38N (2.0 Sv) naturally matches the NADW hydrological properties, as the definition we propose for the NACB is based on a necessary transformation of waters lighter than s1 5 32.20 into denser waters. b. Warm limb circulation north of 538N In order to investigate the way warm waters penetrate the northernmost part of the Atlantic, we extend the

trajectories of the particles intercepted at 538N to new final limits (expt 2b). We use Denmark Strait and the Iceland–Scotland passage to diagnose the entrance into the Arctic Sea and we also intercept the particles that flow back to 538N to describe a first cyclonic rotation in the subpolar gyre. The fraction of SW actually transmitted to the Arctic Sea (Fig. 5a), before any recirculation in the subpolar gyre, uses mostly the Iceland– Scotland passage (1.3 Sv) and less frequently Denmark Strait (0.4 Sv). This latter transport is probably related to the discretization chosen for the Denmark Strait interception line, likely to capture a residual fraction of the subpolar gyre. Of the 3.0 Sv that recirculate back to 538N, a small fraction (roughly 1 Sv) is seen to follow closely the topography and is very sensitive to enhanced mixing with overflow waters. The remaining component follows a smaller scale recirculation path in the subpolar gyre. As discussed in Blanke et al. (1999), one must not forget that the streamfunction we picture corresponds to a time and vertical integration of a four-dimensional transport field. It cannot account for inflows and outflows occurring at the same horizontal grid point but at different months or model levels (as it often happens when the initial and final interception sections of a Lagrangian experiment are identical). Only the dominant movement, in terms of transport, is displayed. As a result, even though the SW inflow documented north of 538N amounts to 4.7 Sv (see Table 5), Fig. 5a does not exhibit streamfunction contours above 4.0 Sv (using a 0.5 Sv contour interval). In a similar way we display the northern pathways for TW (Fig. 5b). A very small fraction of TW that reaches 538N is immediately transmitted to the Arctic Sea, almost exclusively through the Iceland–Scotland passage (0.4 Sv). The rest (0.4 Sv) follows the subpolar gyre before more complex circulations or conversions eventually occur. The pathways obtained for LTW (Fig. 5c) differ slightly, as a dominant fraction (1.8 Sv) is


FIG. 5. Same as Fig. 3 but for waters originating in 538N and transmitted to the Iceland–Scotland passage, to Denmark Strait or recirculating to 538N as computed in expt 2b. The value of the streamfunction is arbitrarily set to 0 above Greenland. The names of the water masses are related to their initial definition and state at 5.38N: (a) SW with a 0.5 Sv contour interval, (b) TW with a 0.1 Sv contour interval, (c) LTW with a 0.25 Sv contour interval, and (d) IW with a 0.25 Sv contour interval.

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FIG. 6. Same as Fig. 3 but for waters flowing from Bering Strait to 5.38N as computed in expt 3. The value of the streamfunction is arbitrarily set to 0 above America. The contour interval is 0.25 Sv. The inset displays the Arctic Sea and shows an anticyclonic circulation in the Beaufort gyre.

directly transmitted to the Arctic Sea, whereas only 1.0 Sv recirculates back to 538N. Equivalent paths (Fig. 5d) are obtained for IW, except with smaller intensities (0.8 and 0.7 Sv, respectively). Before any recirculation within the subpolar gyre, the warm inflow to the Nordic seas amounts to 4.8 Sv (Table 5). After successive rotations in the gyre, some additional 4.2 Sv will reach the Norwegian Sea (3.2 Sv) or the Greenland Sea (1.0 Sv) and explain the annual mean inflows diagnosed over Denmark Strait (1.8 Sv) or the Iceland–Scotland passage (7.2 Sv) sections (Fig. 1). 5. Picturing and quantifying the cold limb of the NACB a. Bering Strait waters Though Bering Strait waters do not belong to the cold limb of the NACB, because they are not related to a modification of Atlantic tropical waters, we find it useful to display their pathways within the domain we study since they contribute to the NADW flow at 5.38N. In this specific Lagrangian experiment (expt 3), initial positions for particles are chosen at the level of Bering Strait with a total transport of 0.8 Sv. All particles are intercepted at 5.38N as UDW (0.3 Sv), MDW (0.4 Sv), and LDW (0.1 Sv), according to the definitions given in Table 2. The pathways corresponding to the flow of Bering Strait waters to 5.38N are shown in Fig. 6. After

some intense anticyclonic recirculation in the Beaufort gyre, the whole flow reaches the subpolar gyre through Denmark Strait in the shape of a boundary current along the eastern coast of Greenland. Some recirculation in the subpolar gyre is then evidenced before a dominant route west of the Mid-Atlantic Ridge (see comments for the UDW flow in the next paragraphs) and an interception at 5.38N within a narrow DWBC. b. NADW circulation south of 538N In a way somewhat equivalent to the analysis of the warm limb of the NACB (expt 2a), we diagnose the way NADW flows from the northern Atlantic to 5.38N. We consider now as initial particles at 5.38N the final positions of particles of experiment 1 belonging to the cold limb of the Atlantic overturning. We integrate the corresponding trajectories backward in time back to the exact same vertical sections defined in experiment 1a. This new Lagrangian experiment (expt 4a) is done separately for the three components of the NADW considered at 5.38N. The upper NADW constituent, UDW, exhibits two different routes of almost equal intensity when flowing from 538N to 5.38N (5.0 Sv; Fig. 7a). The western route (2.5 Sv), west of the Mid-Atlantic Ridge, resembles much the pathways obtained previously for the Bering Strait waters. The other route (2.5 Sv), east of the ridge,

FIG. 7. Same as Fig. 3 but for the cold water masses of the NACB reaching 5.38N and originating in 538N or in the Gulf of Cadiz (98W) as computed in expt 4a. The value of the streamfunction is arbitrarily set to 0 above America. The names of the water masses are related to their definition and state at 5.38N: (a) UDW with a 0.5 Sv contour interval, (b) MDW with a 0.5 Sv contour interval, and (c) LDW with a 0.25 Sv contour interval.

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connects with the DWBC along the coast of America after a westward turn south of the Azores and another crossing of the ridge between 508 and 558W. This circulation corresponds to some circulation schemes inferred for instance for LSW in the east Atlantic (Paillet et al. 1998). Indeed, part of the LSW (the lightest origin of NADW that we can easily associate with UDW) is known to mix with the lower part of Mediterranean Water (MW) in the eastern subtropical Atlantic (Talley and McCartney 1982; Harvey and Arhan 1988; Cunningham and Haine 1995; Paillet et al. 1998). The fraction of UDW that originates in the Gulf of Cadiz (1.9 Sv, Fig. 7a) is likely to include most of the MW outflow at Gibraltar plus a deep-water contribution from the northern Atlantic (538N), flowing close to the Iberian Peninsula and crossing our section at 98W before turning westward toward the DWBC. It is worth noting that the pathways diagnosed from 98W resemble the eastern route diagnosed from 538N, at least for the portion of the flow directed westward: this comforts the hypothesis of LSW mixing with MW in the east Atlantic. Finally, as the total amount of UDW involved in the cold limb is 9.8 Sv, we diagnose 2.9 Sv of UDW being formed (by conversion or entrainment of warmer waters) within the subtropical Atlantic (south of 538N), likely within the southern portion of the subpolar gyre. Figure 7a confirms the necessary existence of a front between the upper layers of the MW outflow and the southward export of subpolar UDW, at the latitude of the Iberian Peninsula. This front was indeed diagnosed by Talley and McCartney (1982) and Paillet et al. (1998) when they addressed the spreading of LSW in the eastern North Atlantic. The MDW flow (Fig. 7b) exhibits mostly a western route from 538N to 5.38N (4.7 Sv), likely because of an enhanced topographic constraint of the ridge on this deeper NADW component. The flow from 98W (0.3 Sv), crosses the Atlantic westward and connects to the DWBC slightly more south (east of the West Indies) than the UDW component. The weak contribution to MDW of waters from the Gulf of Cadiz proves that the outflow at Gibraltar does not include waters deeper (denser) than the lightest NADW component (UDW). We evaluate the amount of MDW being formed south of 538N (i.e., unexplained by the 538N and 98W origins) to 0.6 Sv. Finally, the LDW flow to 5.38N originates almost only in the subpolar gyre (absence of 98W origins, and only 0.1 Sv formed south of 538N, likely in the southern part of the subpolar gyre), amounts to 1.3 Sv (Fig. 7c) and takes dominantly a western route. c. NADW circulation north of 538N We now extend the backward trajectories of experiment 4a that were reaching 538N to new geographical limits, particularly relevant for studying the respective contributions of ISOW and DSOW origins, or the recirculation in the subpolar gyre (expt 4b). For the spe-

TABLE 7. Northern origins of the NADW as defined at 5.38N (in terms of UDW, MDW, and LDW components) and considered at 538N (expt 4b). All transports are expressed in Sv. The particles used in this experiment (and accounting for 11 Sv) are those of expt 4a intercepted at 538N in their backward integration from 5.38N. The UDW origins are investigated separately with respect to the route subsequently followed south of 538N (west or east of 338W). Iceland– Recircu- Southward Denmark Scotland lation flow Strait Passage from 538N (538N) UDW (west of 338W) UDW (east of 338W) UDW MDW LDW Total

0.5 0.0 0.5 1.8 0.5 2.8

0.2 0.6 0.8 0.4 0.1 1.3

1.8 1.9 3.7 2.5 0.7 6.9

2.5 2.5 5.0 4.7 1.3 11.0

cific study of the UDW circulation, we distinguish the particles initially located west or east of 338W at 538N in order to elucidate more easily the origins in the model of the western and the eastern routes. This distinction is no longer necessary for the MDW and LDW circulations because both exhibit almost exclusively western routes to 5.38N, south of 538N. Table 7 sums up the various Arctic origins found for all three NADW components (as defined at 5.38N in terms of potential density). The total southward flow we document at 538N, as the fraction of the NADW related to the cold limb of the NACB, amounts to 11.0 Sv and is made of UDW (5.0 Sv), MDW (4.7 Sv), and LDW (1.3 Sv). The main origin found for this transport is the subpolar gyre itself as more than 60% of the transport (6.9 Sv) is seen recirculating from 538N. Integrating further backward the trajectories that describe this recirculation would certainly lead again to a large fraction still originating in the gyre and additional contributions from the Greenland and Norwegian Seas through Denmark Strait and the Iceland–Scotland passage. For sake of simplicity, we chose not to document explicitly these multiple passages in the gyre, and we only diagnose here the interceptions obtained within the last passage before the southward export, east or west of the Mid-Atlantic Ridge. The contribution from Denmark Strait (2.8 Sv) is more than double that of the Iceland–Scotland passage (1.3 Sv). Most of the flow feeds the MDW component (1.8 Sv), and the UDW and LDW layers contribute equally (0.5 Sv). It is worth noting that the main LDW origins are found in Denmark Strait, consistent with the idea that the densest NADW enters the western North Atlantic as DSOW. On the contrary the lighter ISOW corresponds to the upper levels (UDW and MDW) of the NADW as defined at 5.38N. As DSOW and ISOW are likely to mix together, and also with LSW within the Labrador Sea and with MW within the eastern Atlantic, it seems reasonable not to obtain an exact correspondence between LDW and DSOW and between

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MDW and ISOW. The weakness of the LDW contribution (1.4 Sv) to the NACB (see Table 3), as diagnosed at 5.38N, may be partially related to a progressive conversion (by turbulent mixing) of the densest components of the Arctic overflow waters into lighter MDW within the DWBC. The intense contribution of Denmark Strait to the MDW flow (1.8 Sv) must include water that first flows within the deepest layers of the North Atlantic, as DSOW, before being slowly upwelled and transformed into lighter waters when reaching the equatorial domain (Dickson and Brown 1994). The direct transmission of Nordic seas waters to the equator (excluding Bering Strait waters) only amounts to 4.1 Sv. As the total outflow of ISOW and DSOW is 9.8 Sv (6.1 Sv and 3.7 Sv at Denmark Strait and Iceland–Scotland passage, respectively; see Fig. 1) 5.8 Sv are composed of waters with a Pacific origin (0.8 Sv) and of NACB waters that must recirculate one or several times within the subpolar gyre before joining the tropics. The pathways obtained for waters originating in Denmark Strait are very similar for all NADW components (Fig. 8a), with a southwestward flow along the coast of Greenland, before an incursion within the Labrador Sea that appears all the more pronounced as it affects the uppermost levels of the flow (UDW). Pathways from the Iceland–Scotland passage (Fig. 8b) first follow the eastern flank of the Reykjanes Ridge before a bifurcation occurs between a part that turns westward to the Irminger Basin and another fraction that proceeds southward to 538N in the eastern basin. The topographic steering of the ISOW to the Irminger Basin is a better-known circulation feature than the southward extension (e.g., Harvey and Theodorou 1986), yet the latter was detected to about 458N by Lee and Ellett (1965). The recirculating waters of the subpolar gyre appear constrained by the topography too (Fig. 8c). The Lagrangian approach lets us diagnose a triple origin for the waters that will eventually feed the UDW eastern route, south of 538N (Fig. 9, marks A, B, and C). We find first (A) 0.6 Sv flowing from the Norwegian Sea into the Iceland Sea (i.e., west of the Rockall Plateau) through the Faroe Bank Channel (the southernmost fraction of the Iceland– Scotland passage) and joining the UDW east of the MidAtlantic Ridge (without turning northeastward along the northern edge of the Reykjanes Ridge). Another fraction of the flow is made of recirculating waters around 538N, either within a cyclonic cell around the Rockall Plateau (B) or within an anticyclonic circulation bringing waters from the west (C). The first branch (B) conveys NAC water mixed in the vicinity of the Faroe Bank Channel with ISOW and entrained at deeper (UDW) levels. The latter (C) is likely related to the main path of the LSW circulation toward the eastern basin (Paillet et al. 1998). 6. Discussion We can diagnose the relative contribution of the northern Atlantic (subpolar gyre and Arctic seas) and

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Mediterranean outflow (water mass formation within the Mediterranean Sea and entrainment within the Gulf of Cadiz) by testing the individual trajectories of the particles contributing to the NACB (16.8 Sv). We ran another experiment (expt 5) keeping the same initial particles as in experiment 2a but removing the interception sections at 538N and 98W. We only define flags to check whether one particle crosses either or both sections before returning to 5.38N. The sequence of possible multiple crossings is recorded so that a diagnostic of the first and last interceptions can be made. The results are summed up in Table 8. They distinguish the fractions of the NACB that involve only the Atlantic north of 538N or the Atlantic east of 98W (and the associated Mediterranean Sea) from the overturning that confines itself to the subtropical Atlantic or from more complex behaviors that imply movements inside both northern and eastern subdomains. These latter particles are then classified according to the first and last sections (538N or 98W) they visit along their pathways (from and back to 5.38N). We find that roughly half of the NACB (i.e., 8.6 Sv) corresponds to trajectories that extend north of 538N without penetrating to the Gulf of Cadiz. This is the fraction of the NACB purely associated with deepwater formation in the subpolar gyre and Arctic seas, with likely direct contributions from Denmark Strait and Iceland–Scotland passage overflows. On the other hand we obtain 1.6 Sv of the NACB confined in the eastern Atlantic: this transport is made of pure Mediterranean Water and of North Atlantic Central Water being entrained within the eastern Gulf of Cadiz. This result is in reasonable agreement with estimates of the spreading of the Mediterranean outflow [e.g., 1.9 Sv at 78W in Baringer and Price (1997), or 2.2 Sv at 88W in Ochoa and Bray (1991)]. A large fraction of the NACB involves combined incursions north of 538N and east of 98W and amounts to 3.1 Sv. However most of the particles that explain this circuit finish their overturning in the North Atlantic by a passage in the subpolar gyre (2.5 Sv) and are likely to acquire physical properties characteristic of NADW. The fraction of the NACB that journeys north of 538N amounts to 11.7 Sv, which only underestimates slightly the 14 Sv proposed by McCartney and Talley (1984). Finally, 21% of the NACB (3.5 Sv) consists of an Atlantic overturning confined south of 538N and west of 98W. Figure 10 shows the positions of these particles where they switch from a ‘‘warm’’ state (s1 , 32.20) to a ‘‘cold’’ state (s1 . 32.20). Multiple positions may be obtained for one single particle when successive transitions between the warm and cold NACB limbs occur along the same trajectory. The mass transport carried by each particle is then divided equally among these positions. A cluster analysis is applied to the resulting discrete transport field in order to highlight a few individual regions of interest. Three clusters of equal significance are labeled on Fig. 10 and explain roughly 80% (i.e., 2.8 Sv) of the subtropical overturning. The

FIG. 8. Same as Fig. 3 but for all NADW components (UDW, MDW, and LDW, as defined at 5.38N) reaching 538N southward, as computed in expt 4b. The value of the streamfunction is arbitrarily set to 0 above Greenland: (a) for a Denmark Strait origin with a 0.5 Sv contour interval, (b) for an Iceland–Scotland passage origin with a 0.5 Sv, and (c) for recirculation around 538N with a 0.5 Sv contour interval.

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FIG. 9. Same as Fig. 8 but for UDW originating in the Iceland–Scotland passage or recirculating from 538N, and joining the UDW eastern route (east of the Mid-Atlantic Ridge). The contour interval is 0.25 Sv. Marks A, B, and S denote three distinct origins discussed in the text.

knowledge of the initial and final hydrographical states at 5.38N for the involved particles lets us conclude that the northernmost cluster is linked to a transformation of SW into UDW. On the contrary the two other clusters are mostly linked to IW conversion within the subtropical Atlantic. We compute the vertical mass streamfunction deduced from the zonal integration of the 3D movement of the particles of each cluster (Fig. 11). In this figure, contours now document the conversion on a latitude–depth plane and can be used to evidence the characteristics of the inflow and outflow at 5.38N. Figure 11a displays a dominant conversion of IW into UDW south of 158N at the level of the vertical shear between the northward Caribbean Current and the equatorward western boundary undercurrent, as well as some conversion of lighter water masses within the southernmost fraction of the subtropical gyre. Note that McCartney (1993) also diagnosed strong diapycnal exchanges between the western boundary counterflows of IW and UDW. A more remote conversion is found for the second cluster (Fig. 11b): the transformation of IW into denser water masses occurs in the eastern Atlantic at the latitude of the specified cluster (208 to 308N). The dominant feature of the overturning linked to this cluster is a cyclonic circulation (not shown), south of the subtropical gyre, that first exports IW eastward from the northwestward flowing boundary current, and then redirects the converted flow westward, at a deeper level and at

the latitude of UDW originating at 98W, in a manner somewhat similar to Reid’s (1994) scheme. We note that the water mass conversion in this second cluster occurs at latitude where IW loses its salinity minimum signature in the interior North Atlantic (Tsuchiya 1989). The last cluster, north of the Azores, involves exclusively SW (Fig. 11c) and the conversion occurs in two stages. Intense recirculation (not shown) within the subtropical gyre first injects SW at deeper levels than the initial flow at 5.38N. Entrainment with southward flowing UDW is then achieved north of 408N, before transmission to 5.38N. 7. Sources of error a. OGCM-based errors Our study relies on a careful analysis of the tracer and velocity fields of an OGCM. Though the physics of the global model is fairly sophisticated, several sources of error are likely to impair the overall realism of the results. First, this study deals with a coarse resolution simulation, both in space (model mesh) and time (climatological forcing, frequency of the model output). The model represents ocean eddies very poorly. The vertical resolution in layers deeper than the main thermocline is another source of inaccuracy. The weakness of the BW flow north of the equator (see Fig. 2) is an

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TABLE 8. Relative contributions of the northern Atlantic and the Mediterranean outflow to the NACB as deduced from expt 5 by marking NACB particles when they reach 538N or 98W. All transports are expressed in Sv. With incursion north of 538N

No incur98W 538N sion north last last Total of 538N

With incursion east of 98W 98W first 538N first Total No incursion east of 98W

0.0 0.6 0.6

1.8 0.7 2.5 8.6

1.8 1.3 3.1

1.6 3.5

obvious sign of the lack of resolution available in the bottom layers. Beckmann and Do¨scher’s (1997) bottom boundary layer parameterization and the use of an isopycnal scheme for lateral mixing help the model to deal with overflows. However, no real effort was made in this simulation to constrain the intensity of the mass exchanges through the narrowest straits, by tuning for instance the friction along model lateral boundaries. As a matter of fact the Mediterranean outflow at Gibraltar is more than twice the observed average value (roughly 1 Sv). Fortunately, the relaxation to Levitus (1982) climatology has a major role in reducing the sensitivity of the model to imperfect parameterizations by constraining the interior ocean flow to a reasonable geostrophic

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circulation. The situation is somewhat similar for western boundary currents. The model, with its coarse lateral resolution, struggles in representing sharp and intense currents. Nevertheless, the prescription of the interior ocean return flow (mostly in terms of a Sverdrup balance), in addition to mass conservation, allows the recovery of fairly appropriate mass transports along the western boundary. Model and method deficiencies can define limits to the scope of our results, but the overall consistency found with previous analyses make them credible. Throughout this study we introduced several comparisons with independent mass transport estimates based on observations. The variability of the Eulerian transport of the major currents in the equatorial Atlantic (Equatorial Undercurrent, NECC, etc.) was already investigated (Blanke et al. 1999) and was shown to match fairly well some available observations, despite internal restoring to climatologies within the tropical domain. Then, for the same simulation, additional comparisons with observed flows through critical sections have been discussed on a global scale (Blanke et al. 2001). We cannot say we trust the model velocity everywhere, but we try to compare model mass transport estimates with published circulation schemes, whenever and wherever it is possible, to uncover large discrepancies or fair agreements.

FIG. 10. Intersection locations of the NACB with the s1 5 32.20 isopycnal. Calculations are done for the fraction of the NACB confined south of 538N and west of 98W as computed in expt 5. Only positions for switches from a ‘‘warm’’ state (s1 , 32.20) to a ‘‘cold’’ state (s1 . 32.20) are shown. A cluster analysis applied to the ensemble of positions (weighted by the transport they account for) helps locate three specific regions that are discussed in the text.

FIG. 11. Vertical streamfunction [see Blanke et al. (1999) for details about the calculation] related to the zonally integrated transport of selected particles as computed in expt 5 and displayed as clusters in Fig. 10: (a) for cluster ‘‘1,’’ (b) for cluster ‘‘2,’’ and (c) for cluster ‘‘3’’. The value of the streamfunction is arbitrarily set to 0 at the ocean bottom. Negative values are dotted. The contour interval is 0.1 Sv.

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b. Lagrangian analysis–based errors Temperature (and salinity) relaxation to climatological values induces sinks and sources of heat (and salt) in the interior ocean. In the framework of a Lagrangian analysis, this restoring presents one major drawback as, along any given three-dimensional trajectory, temperature (and salinity) may exhibit jumps associated with these sources and sinks. This is problematic in the sense that one would rather associate Lagrangian changes in tracers to the true and direct effect of turbulent mixing (accounted for mostly in the model by both isopycnal and vertical mixing schemes). However, one can also consider this restoring as part of the model physics since it tends to mimic the effects of poorly performing subgrid-scale parameterizations. It involves a nonlocal redistribution of heat and salt, unlike lateral and vertical mixing that conserve heat and salinity, through exchanges between adjacent gridcells. As the simulation is equilibrated (there is no substantial drift after over 10 years of integration), the internal sources and sinks of heat and salt must balance exactly the surface heat and evaporation-minus-precipitation fluxes. We could check that the annual mean of the atmospheric fluxes is almost zero when integrated over the global domain. Therefore, the relaxation does not create heat or salt on average. One other major problem related to Lagrangian tracing techniques of selected water masses is linked to the time scales of advection processes. Lagrangian trajectories must be integrated long enough to document the full extent of basin-scale or global-scale water mass movements (several hundreds or even thousands of years), whereas direct OGCM simulations seldom exceed a few tens of years because of CPU limitations or inherent drifting problems. One way to bypass this issue is the development of offline Lagrangian calculations (as performed in this study), making use of archived velocity and tracer fields from a nondrifting simulation, and relying on multiple loops over the archived period. We are currently investigating the sensitivity of Lagrangian trajectories in the North Atlantic to the sampling of the velocity dataset. One main result is the convergence of trajectory error measurements for sampling time intervals shorter than the dominant period of forced and internal variability, that is, in the case of this noneddy resolving model, the monthly definition of the surface and internal forcing functions. Larger samplings not resolving the seasonal cycle fail to recover accurate Lagrangian calculations. Once an appropriate sampling period for the OGCM direct simulation is selected, the error of the Lagrangian analysis is conveniently addressed by means of equivalent backward and forward trajectory calculations for any diagnosed directional transport (see section 2b). The value of the prescribed maximum transport allocated to each particle can be chosen small enough to certify that the resulting difference on independent backward and

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forward estimates of the same mass transfer is smaller than a given acceptable error. All Lagrangian computations discussed in this paper could be run both independent ways. The size of the resulting error is roughly of the order of the infinitesimal transport given to each individual particle, here 10 22 Sv per documented month. It allows us to define all our Lagrangian transports with an accuracy better than 0.1 Sv. 8. Conclusions We used three-dimensional monthly velocity fields from a global OGCM to picture and analyze the conversion of warm water entering the North Atlantic at 5.38N with potential densities lighter than s1 5 32.200 and leaving the North Atlantic at the same latitude as cold waters with potential densities above s1 5 32.20. The description of this conversion by means of individual trajectories of numerical particles allows us to link in a dynamic way both flows and therefore to define a warm limb and a cold limb explained by the very same set of particles. Careful analysis and comparison of the hydrographical properties at the entrance and exit of the North Atlantic led us to an accurate estimate of water mass conversions occurring north of 5.38N, with respect to the classification proposed by Lux et al. (2001). The Lagrangian analysis ensured an adequate distinction between meanders of the large-scale circulation or conversions occurring within the same limb of the Atlantic overturning, and true water mass modifications related to journeys in the Mediterranean Sea (or immediate neighborhood) or in the Arctic Ocean. We focused on the particles explaining these latter modifications and built circulation schemes, first for the warm limb of the NACB from 5.38N to the Gulf of Cadiz (98W) and 538N and beyond, then for the cold limb from the Arctic straits to 538N and then from 538N or 98W to 5.38N. This study corresponds to a Lagrangian interpretation of the velocity and the tracers calculated by an OGCM. Several sources of error, chargeable either to the OGCM or to the Lagrangian analysis, were identified and discussed in the previous section. The mass transfers that we compute are obtained from a simulation constrained in the ocean interior to observed climatologies for temperature and salinity (Levitus 1982). Appropriate transports are adjusted by the model along coastlines and within the surface mixed layer where the constraint is relaxed to balance the interior circulation. As a result there is an overall fair conformity of the modeled circulation with formerly published schemes and transport estimates. Besides different known flow patterns, our results show the relative importance of several pathways with respect to others, depending on the considered water mass. For the NACB warm limb in the subequatorial region we can contrast a purely western boundary route and a more eastern inflow as in Bourles et al. (1999),

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and in the subtropical Atlantic a Caribbean Sea route and a more direct path north of the West Indies, as already shown by Schmitz and McCartney (1993). For the cold limb, we diagnose the origins of the eastern route (east of the Middle Atlantic Ridge) for the southward flowing UDW and relate them to LSW flowing from the western Atlantic, as suggested by Paillet et al. (1998), and to a contribution from the Norwegian Sea through the Iceland–Scotland passage. The relative contributions of Denmark Strait and the Iceland–Scotland passage to the direct formation of the densest NADW layers without further recirculation in the subpolar gyre were investigated too. They amount to 2.8 Sv and 1.3 Sv, respectively. These circulation patterns do not contradict some known schemes for the North Atlantic (Schmitz and McCartney 1993; Dickson and Brown 1994) and may help the interpretation of the results from other models or from direct or indirect tracer measurements. From the importance granted to the interannual or interdecadal variability of the NADW formation rate in the North Atlantic with respect to its impact on the European or even global climate, we wish to develop the same kind of diagnostics applied to contrasted climatological states of the North Atlantic circulation, as those produced in climate coupled circulation models running greenhouse scenarios. An accurate Lagrangian documentation of the large-scale circulation pattern would indeed provide valuable help in understanding the tremendous changes in heat distribution caused by such drastic climate changes. Other directions of investigation include the development of Lagrangian diagnostics suitable for the analysis of long ocean simulations that include interannual variability, so as to trace in the deep ocean water mass formation anomalies, as those documented for instance for LSW. Acknowledgments. We wish to thank Gurvan Madec and all our TRACMASS partners for useful discussions. It is a pleasure to acknowledge comments and suggestions by two anonymous reviewers and by the editor, Lynne D. Talley, which considerably improved the legibility of the paper. Support for this study has been provided by the Centre National de la Recherche Scientifique (CNRS) for BB, the Institut Français de Recherche pour l’Exploitation de la Mer (IFREMER) for MA, the Universite´ de Bretagne Occidentale for SS, and the Ministe`re de l’E´ducation Nationale, de la Science et de la Recherche for KP. This work is also supported by a grant from the European Community (TRACMASS project, Contract MAS3-CT97/0142). Lagrangian calculations were performed with the computational resources available at LPO, at the Centre de Brest of IFREMER, and at the CNRS Institut du De´veloppement et des Ressources en Informatique Scientifique.

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REFERENCES Arhan, M., A. Colin de Verdie`re, and L. Me´mery, 1994: The eastern boundary of the subtropical North Atlantic. J. Phys. Oceanogr., 24, 1295–1316. Baringer, M. O., and J. F. Price, 1997: Mixing and spreading of the Mediterranean outflow. J. Phys. Oceanogr., 27, 1654–1677. Beckmann, A., and R. Do¨scher, 1997: A method for improved representation of dense water spreading over topography in geopotential-coordinate models. J. Phys. Oceanogr., 27, 581–591. Blanke, B., and P. Delecluse, 1993: Variability of the tropical Atlantic Ocean simulated by a general circulation model with two different mixed-layer physics. J. Phys. Oceanogr., 23, 1363–1388. ——, and S. Raynaud, 1997: Kinematics of the Pacific Equatorial Undercurrent: An Eulerian and Lagrangian approach from GCM results. J. Phys. Oceanogr., 27, 1038–1053. ——, M. Arhan, G. Madec, and S. Roche, 1999: Warm water paths in the equatorial Atlantic as diagnosed with a general circulation model. J. Phys. Oceanogr., 29, 2753–2768. ——, S. Speich, G. Madec, and K. Doö¨s, 2001: A global diagnostic of interocean mass transfers. J. Phys. Oceanogr., 31, 1623–1632. Bo¨ning, C. W., and P. Herrmann, 1994: Annual cycle of poleward heat transport in the ocean: Results from high-resolution modeling of the North and equatorial Atlantic. J. Phys. Oceanogr., 24, 91–107. Bourles, B., R. L. Molinari, E. Johns, W. D. Wilson, and K. D. Leaman, 1999: Upper layer currents in the western tropical North Atlantic (1989–1991). J. Geophys. Res., 104, 1361–1375. Broecker, W. S., 1991: The great ocean conveyor. Oceanography, 4, 79–89. Cunningham, S. A., and T. W. N. Haine, 1995: Labrador Sea Water in the eastern North Atlantic. Part I: A synoptic circulation inferred from a minimum in potential vorticity. J. Phys. Oceanogr., 25, 649–665. Curry, R. G., M. S. McCartney, and T. M. Joyce, 1998: Oceanic transport of subpolar climate signals to mid-depth subtropical waters. Nature, 391, 575–577. Dickson, R. R., and J. Brown, 1994: The production of North Atlantic Deep Water: Sources, rates, and pathways. J. Geophys. Res., 99, 12 319–12 341. ——, E. M. Gmitrowicz, and A. J. Watson, 1990: Deep-water renewal in the northern North Atlantic. Nature, 344, 848–850. Doö¨s, K., 1995: Interocean exchange of water masses. J. Geophys. Res., 100, 13 499–13 514. Fine, R. A., 1995: Tracers, time scales, and the thermohaline circulation. The lower limb in the North Atlantic Ocean. Rev. Geophys. 33 (Suppl.: U.S. National Report to International Union of Geodesy and Geophysics 1991–1994),E 1353–1365. Friedrichs, M. A. M., and M. Hall, 1993: Deep circulation in the tropical North Atlantic. J. Mar. Res., 51, 697–736. Gordon, A. L., 1986: Interocean exchange of thermocline water. J. Geophys. Res., 91, 5037–5046. Harvey, J. G., and A. Theodorou, 1986: The circulation of Norwegian Sea overflow water in the eastern North Atlantic. Oceanol. Acta, 9, 393–402. ——, and M. Arhan, 1988: The water masses of the central North Atlantic in 1983–84. J. Phys. Oceanogr., 18, 1855–1875. Jackett, D. R., and T. J. McDougall, 1995: Minimal adjustment of hydrographic data to achieve static stability. J. Atmos. Oceanic Technol., 12, 381–389. Krauss, W., 1986: The North Atlantic Current. J. Geophys. Res., 91, 5061–5074. ——, 1995: Currents and mixing in the Irminger Sea and in the Iceland Basin. J. Geophys. Res., 100, 10 851–10 871. Lee, A., and D. Ellett, 1965: On the contribution of the overflow water from the Norwegian Sea to the hydrographic structure of the North Atlantic Ocean. Deep-Sea Res., 12, 129–142. Levitus, S., 1982: Climatological Atlas of the World Ocean. NOAA Prof. Paper 13, 173 pp. and 17 microfiche. Lux, M., H. Mercier, and M. Arhan, 2001: Interhemispheric exchang-

MAY 2002

BLANKE ET AL.

es of mass and heat in the Atlantic Ocean in January-March 1993. Deep-Sea Res., 48, 605–638. Madec, G., and M. Imbard, 1996: A global ocean mesh to overcome the North Pole singularity. Climate Dyn., 12, 381–388. ——, P. Delecluse, M. Imbard, and C. Le´vy, 1998: OPA 8.1 Ocean General Circulation Model reference manual. Notes du Poˆle de Mode´lisation de l’Institut Pierre-Simon Laplace, Vol. 11, 91 pp. [Available from LODYC, Universite´ Pierre et Marie Curie, Case 100, 4 place Jussieu, 75252 Paris Cedex 05, France.] MacDonald, A. M., and C. Wunsch, 1996: An estimate of global ocean circulation and heat fluxes. Nature, 382, 436–439. McCartney, M. S., 1993: Crossing of the equator by the deep western boundary current in the western Atlantic. J. Phys. Oceanogr., 23, 1953–1974. ——, and L. D. Talley, 1984: Warm-to-cold water conversion in the northern North Atlantic Ocean. J. Phys. Oceanogr., 14, 922– 935. Molinari, R. L., R. A. Fine, and E. Johns, 1992: The deep western boundary current in the western tropical North Atlantic Ocean. Deep-Sea Res., 39, 1967–1984. Ochoa, J., and N. A. Bray, 1991: Water mass exchange in the Gulf of Cadiz. Deep-Sea Res., 38 (Suppl.), S465–S503. Paillet, J., M. Arhan, and M. S. McCartney, 1998: Spreading of Labrador Sea Water in the eastern North Atlantic. J. Geophys. Res., 103, 10 223–10 239. Pickart, R. S., 1992: Water mass components of the North Atlantic deep western boundary current. Deep-Sea Res., 39, 1553–1572. Rahmstorf, S., 1997: Risk of sea-change in the Atlantic. Nature, 388, 825–826. Reid, J. L., 1994: On the total geostrophic circulation of the North Atlantic Ocean: Flows patterns, tracers, and transport. Progress in Oceanography, Vol. 33, Pergamon, 1–92. Rhein, M., 1994: The deep western boundary current: Tracers and velocities. Deep-Sea Res., 41, 263–281. ——, L. Stramma, and U. Send, 1995: The Atlantic Deep Western

1451

Boundary Current: Water masses and transport near the equator. J. Geophys. Res., 100, 2441–2457. Rintoul, S. R., 1991: South Atlantic interbasin exchange. J. Geophys. Res., 96, 2675–2692. Schmitz, W. J., Jr., 1996: On the World Ocean circulation: Volumes 1 and 2. Woods Hole Oceanogr. Inst. Tech. Rep. WHOI-96-03, 391 pp. ——, and M. S. McCartney, 1993: On the North Atlantic circulation. Rev. Geophys., 31, 29–49. Sloyan, B. M., and S. R. Rintoul, 2000: Estimates of area averaged diapycnal fluxes from basin scale budgets. J. Phys. Oceanogr., 30, 2320–2341. Smethie, W. M., Jr., 1993: Tracing the thermohaline circulation in the western North Atlantic using chlorofluorocarbons. Progress in Oceanography, Vol. 31, Pergamon, 51–99. ——, R. A. Fine, A. Putzka, and E. P. Jones, 2000: Tracing the flow of North Atlantic Deep Water using chlorofluorocarbons. J. Geophys. Res., 105, 14 297–14 323. Smith, W. H. F., and D. T. Sandwell, 1997: Global sea floor topography from satellite altimetry and ship depth sounding. Science, 277, 1956–1962. Speich, S., B. Blanke, and G. Madec, 2001: Warm and cold water paths of a GCM thermohaline conveyor belt. Geophys. Res. Lett., 28, 311–314. Stocker, T. F., and A. Schmittner, 1997: Influence of carbon dioxide rates on the stability of the thermohaline circulation. Nature, 388, 862–865. Sy, A., 1988: Investigation of the large-scale circulation patterns in the central North Atlantic: The North Atlantic Current, the Azores Current, and the Mediterranean Water plume in the area of the Mid-Atlantic Ridge. Deep-Sea Res., 35, 383–413. Talley, L. D., and M. S. McCartney, 1982: Distribution and circulation of Labrador Sea Water. J. Phys. Oceanogr., 12, 1189–1205. Tsuchiya, M., 1989: Circulation of the Antarctic Intermediate Water in the North Atlantic Ocean. J. Mar. Res., 47, 747–755.

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