3D structure, mesoscale interactions and potential vorticity conservation in a swoddy in the Bay of...

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Journal of Marine Systems 46 (2004) 47–68

3D structure, mesoscale interactions and potential vorticity

conservation in a swoddy in the Bay of Biscay

Ricardo Sanchez*, Julio Gil

Instuto Espanol de Oceanografıa, Centro Oceanografico de Santander, Promontorio de San Martın s/n. E-39004,

P.O. Box 240 Santander, Spain

Received 16 January 2003; accepted 10 October 2003

Abstract

From an intensive oceanographic survey carried out in the Bay of Biscay (Northern Iberian Peninsula) in August 1998, an

anticyclonic Slope Water Oceanic eDDY (swoddy) was identified and studied. Satellite data showed a central anticyclonic

structure to which two smaller scale cyclones were attached. Its hydrographic parameters showed a central homogeneous core

warmer and saltier than the surroundings situated between 80 and 200 dbar with typical near-constant salinity (35.70) and a

weak potential temperature gradient from 12.75 to 12.55 jC. The dynamic front between the swoddy and its surroundings

appeared intensified and deflected as a response to the interference of the swoddy with the pair of cyclones to the north and

south. This interaction gave rise to E–W asymmetry. Baroclinicity associated with the cyclones caused the most intense

advections of relative geostrophic vorticity and subsequent departures of the isopycnic potential vorticity from the Sverdrupian

potential vorticity. Associated with these anomalies, vertical movements were inferred along the swoddy periphery.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Swoddy; Eddy; Bay of Biscay; Conservation of potential vorticity; Mesoscale

1. Introduction cm/s) slope-intensified flow that advects subtropical

In recent years, a subsurface poleward flow has

been evidenced along the shelf regions of the western

European margin (Ambar et al., 1986; Frouin et al.,

1990; Haynes and Barton, 1990; Pingree and Le

Cann, 1990; Pingree, 1994). This undercurrent, also

known as the Iberian Poleward Current (IPC) appears

as a narrow (f 25–40 km), weak (less than f 20

0924-7963/$ - see front matter D 2004 Elsevier B.V. All rights reserved.

doi:10.1016/j.jmarsys.2003.10.002

* Corresponding author. Present address: Centro de Investigac�aodos Ambientes Costeiros e Marinhos (CIACOMAR)—Universidade

do Algarve. Av. 16 Junho s/n. P-8700-311 Olhao, Portugal. Tel.:

+351-289707087; fax: +351-289706972.

E-mail address: [email protected] (R. Sanchez).

water northwards (Haynes and Barton, 1990). In

winter, this flow reaches the surface and has been

traced with thermal advanced very high resolution

radiometer (AVHRR) satellite imagery as far north in

latitude as the Goban Spur (Pingree, 1993). Pingree

and Le Cann have devoted special attention to the IPC

along the Bay of Biscay with the aid of drifters,

moored instruments, satellite and cruise data. In a

series of papers published beginning in 1990, they

observed that slope water could be detached from the

IPC under abrupt topographic steering, injected into

deep-water regions and result in warm anticyclonic

Slope Water Oceanic eddies (swoddies, e.g., Pingree

and Le Cann, 1992a).

R. Sanchez, J. Gil / Journal of Marine Systems 46 (2004) 47–6848

Swoddies are observed as mesoscale (50–100 km

diameter) water lenses animated with anticyclonic

motion and rotation periods of 15–20 days. In summer,

after the establishment of the seasonal thermocline

capping the feature, they can be traced by a cool spot

above the swoddy core. Direct hydrographic measure-

ments revealed this feature as consequence of upward

doming of isopycnals above the thermocline. Swoddy

thermohaline properties are well discussed in Pingree

and Le Cann (1992a,b), Garcıa-Soto et al. (2002) and

Fernandez et al. (submitted). They are characterized by

an upper homogeneous core with constant T-S pairs

resembling the IPC hydrographic features. They have

been observed to migrate westwards at f 2 cm s� 1

over the periods during which their AVHRR signal is

discernible.

Swoddies have been observed to be composed of a

more complicated system than a simple monopole.

Remote sensing studies often present a multi-polar

structure, with two satellite cyclones attached to the

main eddy. They were observed to rotate as a whole in

near solid-body manner. Garcıa-Soto et al. (2002)

observed that cyclone-swoddy interactions could lead

to significant flow accelerations at the confluence zone.

Additionally, from SeaWIFs scans, diabaric mixing

was inferred between the eddies (Garcıa-Soto et al.,

2002). In the absence of major steering currents, these

effects could exert certain influence on evolution and

decay of the swoddy system.

To study the structure, physical dynamics and bio-

logical functioning of a swoddy in the southern Bay of

Biscay the observational program GIGOVI (Organiza-

cion Trofica y flujo de Materiales en GIros anticiclo-

nicos del GOlfo de VIzcaya) was designed (Fernandez

et al., submitted). A mesoscale swoddy-like structure

that had presumably detached from the IPC before

summer 1998 (named AE6 elsewhere) was sampled

with a biological and physical grid in August 1998. Its

characteristics and dynamical implications are pre-

sented here. Special attention is devoted to the doming

of the seasonal pycnocline and the interactions of slope

waters retained within the swoddy with surrounding

structures. The understanding of these processes is

crucial to explain the ecology of these eddies, which

are likely to be significant for the overall functioning of

the pelagic communities of the region. Earlier studies

of Bay of Biscay swoddies were performed with

satellite data, drifting buoys and hydrographic and

ADCP measurements with poor horizontal resolution.

The present study is addressed with a mesoscale-

resolving CTD grid under the frame of Quasi Geo-

strophic (QG) dynamics, and validated with multi-

sensor satellite data. To our knowledge, this is the first

time this approach is undertaken, what represents a step

forward in the knowledge of swoddy dynamics. The

paper is structured as follows: first, estimates of

swoddy drift prior to and over the sampling period

are presented. Then, the density and velocity fields are

described in detail. Subsequently, an analysis of iso-

pycnic potential vorticity is done. Effects of external

intrusions of submesoscale structures on swoddy cir-

culation are evaluated and areas where departures from

geostrophy are relevant to the conservation of QG

potential vorticity are detected. In the end, vertical

forcing is estimated using the QG omega-equation in

the form developed by Hoskins et al. (1978).

2. Materials and methods

Between 6 and 8 August, a well-defined structure

compatible with previous definitions of swoddies was

observed centered at f 6.2W 45.3N. With this infor-

mation, a cruise was carried out from 12 to 31 August

1998 on board the R/V Professor Shtokman in the box

defined by 44.72N, 7.1W, 45.84N and 5.3W in the Bay

of Biscay (Fig. 1). The data presented here correspond

to 65 CTD stations performed from 14 to 22 August

and are distributed on an almost regular 7� 7 grid with

a separation of 18 km between stations. Satellite

altimeter sea level anomaly (SLA) contours for 12/

22/98 are overimposed for visual reference. (The eddy

center SLA is higher than 12 cm and line interval is 2

cm.) The cruise was performed in two main legs

comprised of stations sampled over quasi-synoptic

periods. Leg 1 (dotted stations) was done between 14

and 18, and leg 2 (triangle stations) between 20 and 22

August 1998. Leg 1 provided definition for both the

location of AE6 core and boundaries in the northeastern

area. The core position and surrounding circulation

along the southern area were well defined in leg 2. For

administrative reasons, the northeastern section of the

study area could not be fully sampled. Stations were

sampled from the surface to a maximum depth of 500

dbar using a Sea Bird SBE 25 CTD system. Addition-

ally, six extra CTD casts were performed down to 2000

Fig. 1. Sampling stations, selected transects for mapping thermohaline fields and derived parameters (I–VII). Cruise leg 1 stations are

positioned as black dots and cruise leg 2 stations are triangles. The inset of the figure gives the location of the study area and sampling stations

in the Bay of Biscay. SLA contours for 22/08/1998 are overplotted for reference of AE6 location.

R. Sanchez, J. Gil / Journal of Marine Systems 46 (2004) 47–68 49

dbar with a Neil Brown Mark III CTD attached to a

rosette sampler. Discrete water samples were taken for

calibration analysis. Thermohaline and other derived

parameters are presented at transects I through VII,

indicated in Fig. 1. For further details of the cruise and

general settings, see Fernandez et al. (submitted) and

Garcıa-Soto et al. (2002).

Internal waves/tides are a common feature in the

Bay of Biscay (Pingree and New, 1995). During the

cruise, it was realized that the sampling region was

disturbed in its upper layers by strong short period

internal waves with frequency of f 4 cph, particularly

affecting the layer between 20 and 40 dbar. They

displaced the thermocline and disturbed stratification.

Dynamic height variations of approximately 2 dyn cm

over the uppermost layer were observed (not shown)

that could bias the calculation of the baroclinic veloc-

ity field. Below 65 dbar, no significant variations due

to the internal waves were observed. To overcome

these problems, first, vertical profiles were smoothed.

Additionally, derived fields were preferentially ana-

lyzed on isopycnal surfaces in order to avoid internal

wave noise.

A 2D univariate spatial analysis of station data

using the Successive Correction method proposed by

Bratseth (1986) was used. It included an analytical

filtering of the analyzed field (Pedder, 1993). The

band-pass response of the filtering was centered at a

wavelength of 35 km, imposed by the station sepa-

ration scheme. The short wavelength noise and non-

resolved structures were filtered out. The results were

two gridded matrices with 7.9� 11.1 km2 cells and

vertically interpolated every 10 m (isobaric analysis)

and 0.01 rh units (isopycnic analysis).

The radar altimeter on board TOPEX/Poseidon (T/

P) and ERS satellites provide continuous samples,

unimpeded by cloud (Crawford et al., 2002). The

poor precision in the knowledge of the geoid prohib-

its the use of an absolute reference level for sea

altimeter measurements. Rather, the anomalies with

respect to a temporal average (SLA) are more precise.

SLA measurements of merged ERS +T/P are pro-

duced at 10-day intervals on a 0.2� 0.2j grid (see Le

Traon et al., 1998). Maps for June–September 1998

were retrieved from the AVISO web gateway (http://

www.aviso.cnes.fr).

3. Results and discussion

3.1. Drift

The T/P-ERS SLA data showed AE6 as one of the

most outstanding mesoscale features in the Bay of

http:\\www.aviso.cnes.fr


Biscay in summer 1998 (Fig. 2). Before July, the AE6

SLA signal was weak. From 3 July, a mesoscaled

anticyclone (AE6) was observed anchored in the

center of the Bay of Biscay. The map sequence after

this time shows that AE6 could be seen as a drifting

multi-polar structure with signal between 11 and 15

cm. The AE6 was bounded by at least two cyclonic

waves at its NNW and SE flanks. They have been

labeled NC and SC (for Northern and Southern

Cyclones, respectively) in Fig. 2, ensemble of 12

August 1998. After this time, AE6 appeared com-

pressed to the E, although this could be a result of

Fig. 2. Sequence of Sea Level Anomaly (SLA) 10-day composition maps

colored. Colorbar units are mm. Contour interval is 10 mm. Dates ar

accompanying cyclones are labeled at the map corresponding to 12/08/19

interpolation bias. The AE6 center as inferred by SLA

drifted south-westwards from the initial position at

5.87W 45.50N on 3 July towards the last monitored

position at 6.75jW 45.18jN on 1 September at a

average drift rate of 1.5 km/day, variable between 1.0

and 3.5 km/day. The maximum displacement occurred

in the last week of August, while the cruise leg 2 was

taking place. The NS and SC seemed to drift attached

to the AE6. Evidence for clockwise translation of

these structures around AE6 is also inferred, although

the lack of resolution prevents the accurate determi-

nation. This tripole characteristic has been observed in

from T/P-ERS altimeter along with ground tracks. Positive SLA are

e from 24/05/1998 until 01/09/1998. The swoddy AE6 and the

98. Areas with interpolation errors greater than 20% are blanked.


previous studies of Bay of Biscay swoddies (Pingree

and Le Cann, 1992a). The eddy drift rate could be

supported by AVHRR imagery analyses also. How-

ever the persistent overcast conditions over the Bay of

Biscay during the summer of 1998 prohibited a more

detailed study (Garcıa-Soto et al., 2002). Between 6

and 8 August, the AE6 core could clearly be observed

at 6.2W 45.3N (Fig. 3). On 29 August, it was

observed some 30 km southeastward in a rather

clouded infrared scan (image not shown). Associated

to the southwestward drift, deep thermohaline time

variations were observed. Three CTD casts carried out

at the core location during leg 1 showed an intense

and consistent cooling and freshening trend (� 0.045

and � 0.005 jC day� 1 at 400 dbar) from 14 through

22 August. It was accompanied by a 0.75 dyn cm fall

in dynamic height (not shown). Drifting estimates

were consistent with the studies of Pingree and Le

Cann (1992a).

The planetary h-effect causes eddies to migrate

westward, irrespective of their polarity. The propaga-

tion speed for quasi-geostrophic (QG) vortices is

equal to the phase speed of Rossby waves (e.g.,

Pingree and Le Cann, 1992a). This speed (for a

theoretically isolated structure), accounted for just

Fig. 3. AVHRR (NOAA-14) SST image of the Bay of Biscay on 7 August

water is in dark shades and cool water in light shades. A colorbar (jC)centered near 44.5N 6W.

4% of the inferred displacement. Pingree and Le Cann

(1992a) postulated that most of the difference must be

explained on the basis of interactions with external

agents, namely steering currents. General circulation

in the Bay of Biscay is anticyclonic with average

velocities of O(2) cm�s� 1 (Pingree, 1993), which may

justify the remaining westward translation of the

structure. The association AE6-NC-SC as revealed

by AVHRR and SLA satellite imagery allowed the

inference of interactions among the structures.

3.2. Water mass field

3.2.1. AVHRR structure

In the period of thermal stratification, surface cool-

ing above the core is a feature of Bay of Biscay

swoddies (Pingree and Le Cann, 1992a,b). SST field

on 7 August 1998 at the AE6 location is presented in

Fig. 3. A conspicuous temperature anomaly was evi-

dent as a cool spot ( < 19 jC) centered at f 6.20W

45.30N and surrounded by a warm annular ring with

temperatures f 19–19.5 jC. The shape of the spiral

wake visible in the thermal structure permitted the

inference of anticyclonic rotation. AE6 appeared as a

complex tripolar system associated with two smaller

1998. The selected box correspond with the sampled domain. Warm

is attached for reference. The swoddy was detected as a cool spot


scale cyclones (NC and SC in Fig. 3) attached at its

SW and NW borders. SeaWiFS chlorophyll-a data

presented by Garcıa-Soto et al. (2002) also showed

the same associations.

Water mass properties were charted along transects

I and II for leg 1 approximately across the eddy center

(Fig. 4; see Fig. 1 for location). The AE6 vertical

temperature structure was characterized by a lens-like

anomaly between below the seasonal thermocline

(100–200 dbar) with near homogeneous salinity val-

ues between 35.69 and 35.71 (gray tones in Fig. 4).

The core was associated with sharp upward and

downward doming of isotherms and isopycnals with

respect to the base of the seasonal thermocline. The

25.50 rh surface sloped up from 30 db in the periph-

eral region to the surface at the top of the swoddy

core, as illustrated in Fig. 5a and c. Below the

seasonal thermocline, AE6 featured a marked down-

ward slope of the isopycnals. For instance, the 27.10

rh surface sloped down from 150 to below 300 db

(Fig. 4). Consequently, the stratification appeared

disturbed by the eddy core, imposing larger vertical

ranges between equally spaced isopycnals. Buoyancy

frequency values were decreased down to 25 cph

versus the more stratified surroundings with frequen-

cies of O(36) cph (not shown). AE6 represented a

conspicuous thermohaline anomaly within Bay of

Biscay waters.

Thermohaline properties of the surrounding NC

were charted along transect II (labeled in Fig. 4e). It

was characterized by high near-surface salinities and

downward slope of isohalines from the surface to

f 150 dbar. It resulted in the consequent density

increase north of 45.6jN. Similar pattern was observed

for the SC between 44.8jN and 45.0jN, what broughtabout a steep density increase south of 45jN (not

shown).

Spiciness is a state variable ideally suited for

characterizing water masses at sharp fronts (Flament,

2002). This variable is most sensitive to isopycnal

thermohaline variations and least correlated with the

density field. It is defined as largest for hot (spicy)

and salty water. Spiciness maps for the 25.50 rh

surface are presented in Fig. 5b and d. Low spiciness

characterized the central location, corresponding with

the cool spot observed in satellite imagery, as a

response to the upward lift of the isopycnals. An

annular ring occurred around the central spot, with

spiciness values >3.81 (corresponding with T-S pairs

>19.46 jC and 35.81 at this density level). This ring

was spatially coincident with the (warm) 19–19.5 jCring evident in Fig. 3. Lower spiciness values were

common around the outer periphery. Intense spiciness

fronts were observed between the swoddy-induced

high spiciness anomaly and the surrounding low-

spiciness waters, especially at the W and E borders.

3.2.2. AE6 core properties

A high-resolution vertical profile was plotted at the

AE6 center on leg 1 (6.2W 45.3N) (Fig. 6). The core

featured a vertical salinity profile with constant 35.701

between 80 and 200 dbar. However, potential temper-

ature decreased slightly with depth, yielding a weak

gradient from 12.75 to 12.55 jC. Eddy-atmosphere

exchanges through ventilation of the upper seasonal

thermocline might be considered to be the cause of

upper core warming. This resulted in densities increas-

ing slightly with depth from 26.99 at 80 dbar to 27.03

at 200 dbar. These were the lowest density values

recorded in the sampling domain. Values presented

here differed from other swoddies presented in the

literature. For instance, Pingree and Le Cann (1992a)

reported sustained values of h = 12.95 jC, S = 35.736and rh = 26.97 over a 195-dbar range (from 65 to 260

dbar) for F90a in July 1990.

Pingree (1994) related strong IPC and winter

warming in the Bay of Biscay with swoddy genera-

tion. With satellite and in situ data, he showed that

the winters 1988/1989 and 1989/1990 were seasons

with marked winter warming in the northern Spanish

continental slope by the IPC. For example in January

1989, maximum temperatures of 13.9 and 11.7 jCwere recorded at the surface and at 210 m, respec-

tively. On 22 January 1998, the eastern Cantabrian

Sea showed T-S pairs of 13.8–35.65 and 13.7–35.72

jC at 10 and 210 db, respectively, the highest values

of the 1992–2002 monthly series (Project ‘Studies on

time series of oceanographic data’, managed by the

IEO, pers. comm.). Swoddy formation was favorable

during the winter of 1997/1998, a season in which

the IPC was also well developed. Historical SST

analysis by Garcıa-Soto et al. (2002) supports this

manifestation. This evidence permits to affirm that

the winter poleward flow was strong in 1998 and of

comparable magnitude to the extreme 1988 and 1990

events.

Fig. 4. Cross-section during cruise leg 1 (northern half of the domain) along transect I: (a) potential temperature, (b) salinity and (c) sigma theta.

Idem across transect II: (d) potential temperature, (e) salinity and (f) sigma theta. The central core is highlighted by the shaded contour in the

salinity distribution. The location of the NC is labeled in (e).


Fig. 5. Horizontal distribution of isopycnal (a) pressure and (b) spiciness at rh = 25.50 (the near-surface level) for cruise leg 1. (c) and (d) idem

for cruise leg 2.


Fernandez et al. (submitted) discuss the biological

properties of AE6. Copepod and cheatognath species

typical of coastal and shelf regions were observed

exclusively inside the eddy core. Moreover, the exclu-

sive presence of two species of diatoms (Dyctyocha

fibula and D. speculum) characteristic of the autumn–

winter period in the central Cantabrian Sea inside AE6

reflected the biological features of the water body that

initially originated the structure.

From the above, it can be observed that AE6 was

very likely slope-sourced and it is possible to classify

it as a swoddy in the Pingree and Le Cann sense (e.g.,

Pingree and Le Cann, 1992a). It can be speculated

that AE6 had slope origin and formed during the

previous winter months, when the IPC peaked along

the Cantabrian shelf. Away from the atmospheric

boundary, the IPC hydrographical and biological

properties could be maintained in this relatively

well-isolated structure.

The spiciness field at 130 dbar showed horizontal

gradients between the swoddy and the surroundings

(Fig. 7). A high spiciness central region (>2.15)

corresponding with the high temperature and salinity

values at the swoddy center. Outside, a low-spiciness

fringe surrounded the central core. The spiciness field

mimicked the asymmetrical shape of the SLA. During

leg 1, the lowest spiciness values were found at the

eastern (down to 2.07) and at the northern (down to

2.09) rims. Isospices extended northeastwards, with

intense fronts along both the eastern and northern

borders (see SLA for 08/22/1998). During leg 2, the

swoddy isospices seemed to align along the NE–SW

axis (see SLA of the last week of August 1998), with

low values at the SE (down to 2.06) and eastern

(down to 2.07) fronts. Spininess analysis showed that

the general swoddy pattern was maintained through-

out the water column. Eddy asymmetry was inferred

from all the available data.

In sum, above the thermocline, isopycnal spiciness

(and temperature) decreased towards AE6 center. This

pattern changed sign with depth and below the ther-

mocline isopycnal spiciness was significantly higher

Fig. 6. Hydrographic parameters for the swoddy core.

Fig. 7. Horizontal distribution of isobaric spiciness at 130 dbar

crossing the AE6 core for (a) cruise leg 1 and (b) for cruise leg 2.


at the eddy core. This reversal of baroclinicity is a

common feature of oceanic eddies and occurs between

the seasonal and the main thermocline (Onken, 1990).

A great deal of literature on the hydrography of eddies

has been devoted to clarifying this topic (e.g., Olson,

1980; Kenelly et al., 1985; Joyce and Kennelly, 1985;

Pingree and Le Cann, 1992a). With a 2D adiabatic

frontogenesis model, Onken (1990) addressed the

creation of reversal of baroclinicity from PV consid-

erations. He proposed that reversal of baroclinicity

could be created either during the frontogenesis pro-

cess or by divergences of the radial flow after detach-

ment of the ring from the source jet. In this sense,

direct and precise measurements of the velocity struc-

ture in warm anticyclonic Gulf Stream rings (e.g.,

Joyce and Kennelly, 1985; Kenelly et al., 1985) have

revealed outward radial velocities at the level of the

seasonal thermocline. Onken (1990) interpreted that

the enhancement of frontal baroclinicity and acceler-

ation of the jet must lead to the increase in the

cyclonic and anticyclonic shear vorticities on its

respective flanks. Finally, conservation of PV requires

stretching and shrinking of vortex tubes and a com-

pensating cross-frontal mass flux that should lead to

divergence on the anticyclonic side (i.e., upward

motion) and convergence (i.e., downward motion)

on the cyclonic side. Similar conclusions were drawn

by the modeling study of Smith et al. (1996). Garcıa-

Soto et al. (2002) observed that the confluence F90a-

cyclones brought about significant flow accelerations

around the swoddy core. Unfortunately, no direct

velocity measurements can be presented for the AE6

swoddy but geostrophic calculations showed strength-

ening of the flow between AE6 and the cyclones (see

Section 3.3) that introduced vorticity intensification at

the confluence zone. Fig. 3 showed outward water

excursions off AE6 towards the SC and the NC that

could reveal mass flux divergence at the AE6 center.

Both of these pieces of evidence are in agreement with

the model of Onken (1990).

Fernandez et al. (submitted) have specifically dealt

with the biological and chemical properties across the

cool spot associated with AE6. They have proven that

swoddies are important from the ecological point of

view. These authors related the upraising of isopyc-

nals below the seasonal thermocline with consequent


elevation of the associated nitracline. They also

observed that this feature is common in other types

of mesoscale anticyclonic eddies (cf. references in

Fernandez et al., submitted). According to their dis-

cussion, the biological consequences of the doming

are basically marked by primary production process-

es. These are enhanced when vertical transport of

nutrients flows towards the euphotic zone. As a result,

depth-integrated chlorophyll-a values were signifi-

cantly higher at the AE6 center than in the surround-

ings. Additionally, the swoddy induced sharp

modification of the planktonic community composi-

tion and associated size-structure. Satellite-borne

observations have also permitted the association of

the cool spot with a clear phytoplanktonic surface

signal at the AE6 center (Garcıa-Soto et al., 2002).

3.3. Mesoscale pattern and geostrophic velocity field

A number of papers have addressed the determi-

nation of circulation in mesoscale features under the

frame of quasi-geostrophic dynamics from CTD data

alone (cf. Tintore et al., 1991; Pollard and Regier,

1992; Allen et al., 1994). However, a number of

factors hamper its full application under certain

sampling conditions (Gomis et al., 2001). Among

them, observational uncertainties (e.g., time/space

distribution of observations) are an important source

of diagnosis errors in the definition of derived fields

(like, for instance, the relative vorticity field) in

surveys like the one presented here. These limitations

are minimized by improving the sampling strategy

whereby the most precise determination of the den-

sity front and its curvature changes can be attained.

Unfortunately, a robust level of no motion could

not be attained for all stations. To allow for the

geostrophic computations for the entire sampled do-

main, geostrophic computations (based on the isopyc-

nal potential function (Montgomery, 1937)) were

done on the rh = 27.17 reference level, although it

was known not to be level of no motion (Fig. 8). At

near surface (rh = 26.66) for both legs 1 and 2, the

isopycnal potential function field showed a central

anticyclonic structure caused by the swoddy presence

(Fig. 8a and b). The dynamic height signal was f 2

dyn cm. Altimetry data showed that the SLA at the

eddy core was f 8 cm above the surroundings waters

(Fig. 2). The dynamic height signature of AE6 com-

puted over the 50–1500 db layer was f 8 dyn cm

(not shown). This difference is comparable to the

dynamic height anomaly from in situ CTD data for

other studies (Pingree and Le Cann, 1992a). Hence,

one should be aware that relevant information

concerning the baroclinic field was lost when adopt-

ing this level.

The size of AE6 is mesoscale and is consistent

with the estimate of the Rossby deformation radius

Rd=( gVH)1/2C� 1, where H = depth scale and gV= g

(q2� q1)q2� 1 is the reduced gravity. For H = 900 m,

gV= 0.002 (see the deep density field in Fig. 9a)

and Rd= 13.4 km. Hence, mesoscale structures of

fpRd =f 40–50 km were expected. The circula-

tory scheme around the main gyre appeared to be

deformed by interaction with cyclonic structures that

were inferred from satellite data. Plus, the dynamic

gradients intensified at the NE boundary in strong

coherence with the SLA image of 08/22/1998, al-

though the lack of sampling stations in this quadrant

prohibited the resolution of curvature changes of this

front (Fig. 8). Leg 1 showed flow intensification at

the N and NE boundaries by the confluence of AE6

with a cyclonic region at its northern edge (NC, Fig.

8a). From leg 2 data, an important association with a

cyclone on the S was also conspicuous (SC, Fig. 8b).

This caused flow rectification along the southern

front along which the most intense dynamic gra-

dients were found. Both cyclones had well-defined

thermohaline signatures at every isopycnic level,

although the strongest gradients were observed at

upper levels. As a result of this, the dynamic picture

at 27.02 rh appeared notably smoothed with respect

to that at the upper level (Fig. 8c and d). For

instance, the meander observed in association with

NC at rh = 26.66 is almost non-existent at this

isopycnal level. Similarly, the flow along the S front

gained negative vorticity, suggesting dwindling of

SC relative strength.

To assess the effect on the estimated baroclinic

flow of the use of a reference surface that was not a

level of no motion, an analysis of geostrophic

velocities on a deeper (900 dbar) reference level

was performed along transect I (Fig. 9a) and com-

pared with geostrophic velocities along the same

transect on the shallower level (Fig. 9b). Baroclinic

velocities were generally surface-trapped with dimin-

ishing values with increasing depth. Maximum val-

Fig. 8. Dynamic topography of the isopycnal potential function at rh= 26.66 on the rh = 27.17 isopycnal surface (dyn cm) for (a) leg 1 and

(b) leg 2. (c) and (d) idem at rh= 27.02. AE6 and the surrounding structures have been labeled.


ues along transect I were 14 cm s� 1 (Fig. 9a). These

were underestimated by a factor of 4 when computed

on the 27.17 level instead of on the 900 dbar level

and, indeed, a reference level of 1000 db or

rh = 27.55 would probably provide a more realistic

description. However, the upper-layer baroclinic field

appeared consistent irrespective of the reference

level. In any case, the sense of motion fits with an

anticyclonic eddy in the analyzed domain. Hence,

the suitability of the 27.17 level for upper-layer

dynamic considerations could be partially supported

and the dynamic picture presented in Fig. 8 justified.

The plot of transect II velocities (Fig. 9c) revealed

that the westerly flow deflected by the confluence

with the NC, which accounted for the apparent

acceleration of the flow north of 45.6jN. During

Leg 2 intensification of the southward geostrophic

flow at f 100 dbar occurred by the confluence with

the CS (Fig. 9d).

Assuming the correction of the geostrophic ve-

locities calculated on the shallower (rh = 27.17) level

with respect to the deeper (i.e., 900 dbar) level, we

could estimate mean velocities of O(15) cm s� 1 as

characteristic of the flow around AE6. Considering a

radius of f 30–40 km, the approximate rotation

period can be estimated as f 15–20 days. This

estimation lays within values reported for other

swoddies, for instance the f 18 days rotation period

inferred from satellite imagery for F90a (Pingree and

Le Cann, 1992a) or f 20 days for other open ocean

eddies (e.g., Benzohra and Millot, 1995).

3.4. Potential vorticity

In isentropic analysis under adiabatic non-dissipa-

tive geostrophic conditions Rossby-Ertel potential

vorticity (PV) is a conservative property and can be

used as a tracer of flow circulation (Rossby, 1940;

Fig. 9. Sigma theta and geostrophic velocities (cm�s� 1) referenced on: (a) 900 dbar for transect I. Geostrophic velocities referenced at: (b) 27.17

rh level for transect I, (c) idem for transect II, (d) idem for transect III, (e) idem for transect IV. Density contours are overplotted in (a) for

reference.


Ertel, 1942). From the hydrostatic approximation and

C-plane geometry, the formulation of Rossby-Ertel

PV may be written (Pollard and Regier, 1992):

PV ¼ �g�1f N2ð1þ f f �1 � FÞ ð1Þ

where C is the Coriolis parameter, g is acceleration

due to gravity, N= � gq� 1qz the Brunt-Vaisala

buoyancy frequency, fC� 1 normalized geostrophic

vorticity and F =N� 2(Uz2 +Vz

2) the Froude number.

The latter term is implicitly accounted for in relative

vorticity calculated isentropically or isopycnally (is-

entropic vorticity, IV (Rossby, 1940)), representing a

coordinate adjustment when working at a vertical

pressure coordinate rather than on an isopycnal

surface. Thus,

IV ¼ fdens¼const ¼ f � f F ð2Þ

Hence, the expression for PV calculated isopyc-

nally, or isentropic potential vorticity (IPV) results as

follows:

IPV ¼ f þ IV

Dp

Dqq

ð3Þ

Where Dp is the separation (thickness) in meters of

two isopycnals. For mesoscale or gyre motions, since

fbC and Fb1 in most cases, PV is similar to the


planetary vorticity term (or Sverdrupian potential

vorticity according to Woods, 1985)

SPV ¼ g�1fN 2 ð4Þ

and the PV field mirrors the stratification field on the

C-plane. However, on the lower limit of the meso-

and submesoscale, neither fC� 1 nor F should be

negligible with respect to unit everywhere.

Dynamical processes may be analyzed based on the

knowledge of the PV field (Hoskins et al., 1985). In

the first approximation, fluid particles must move

along isopycnal surfaces and it is interesting to exam-

ine IPV distributions. With the parameters of the

swoddy system, applying a velocity scale U =O(0.1)

m�s� 1 and a typical length scale L=O(100) km,

Rossby number Ro =U/(CL) = 0.01 and motion is

estimated to be at near hydrostatic and geostrophic

equilibrium. Hence, the study of conservation of

potential vorticity may also be done under the quasi-

geostrophic theory using a derivation of the general-

ized QG omega equation without cancellation of terms

in the C-plane (Hoskins et al., 1978) as shown, for

example, by Pollard and Regier (1992), Viudez et al.

(1996) and Allen and Smeed (1996).

3.4.1. PV field

Detailed IPV transects are presented in Fig. 10 (see

location sketch in Fig. 1). The AE6 core appeared as a

low IPV anomaly between 100 and 200 dbar, and

whose values were always below 0.04�10� 9 m� 1 s� 1

(dark tones in Fig. 10). A study of the contribution of

individual terms in Eq. (1) including the normalized

isentropic vorticity (IVC� 1) and the Sverdrupian

potential vorticity (SPV) was performed on the

27.02 isosurface. This isopycnal level was chosen

because it crossed the eddy core at its upper part

(Fig. 10). At this density, surface homogeneous AE6

core waters defined the pattern of stratification, pres-

sure and vorticity (Fig. 11). This surface was within

the pressure range 80–150 dbar (Fig. 11a and e), with

layer thickness values greater than 16 m (i.e., low

stratification, Fig. 11b and f). The structure showed a

rather circular shape, although the isobars appeared

compressed towards the NNW and at the E. The NC

seemed responsible for the NNW deflection of the

isolines during leg 1. The southern AE6 boundary

also appeared distorted during leg 2, as evidenced by

the uplift of the 27.02 surface up to 50 db in association

with the SC (Fig. 11e).

The normalized isentropic vorticity values (IVC� 1)

at 27.02 were generally small due to the generally low

computed geostrophic velocity shears and the large

radii of curvature (Fig. 11c and g). Maximum values

were of O(0.05) throughout the domain. The IVC� 1

term showed the anticyclonic anomaly imposed by the

swoddy presence, with values lower than � 0.04

associated with the AE6 center. During both cruise

legs, the swoddy appeared constrained by a band of

cyclonic IVC� 1, with maximum values associated

with the NS ( + 0.03) in leg 1 and SC ( + 0.05) in leg

2. Additionally, IVC� 1 increases were inferred at the E

border, corresponding with a structure unresolved by

the sampling scheme. A closer examination of the

IVC� 1 maps and the isopycnic potential function at

the 27.02 isosurface revealed significant advections of

normalized geostrophic vorticity by the isentropic flow

at two locations. IVC� 1 advections were inferred

where AE6 interactions with the surrounding structures

could be resolved. Hence, negative IVC� 1 advection

(from 0.01 to 0.03 over f 10 km) was inferred at the

NC during leg 1, while positive IVC� 1 advection of

the same order of magnitude could be observed at the

SC during leg 2.

The resulting horizontal IPV field has also been

mapped at 27.02 (Fig. 11d and h). AE6 was repre-

sented as a low IPV anomaly (below 0.04�10� 9 m� 1

s� 1) that strongly contrasted with the surroundings. In

agreement with previous descriptions, maximum gra-

dients were found at the NW and SE boundaries. IPV

contours almost followed the dynamical fronts or the

stratification field for most of the domain (see Fig.

11b and f in terms of layer thickness). However, at

some sites, both dynamic height and IPV fields were

not parallel to each other, at least at the level of the

AE6 core. The source of these discrepancies seemed

to be imposed by flow deflections associated with the

NC and SC. As with the IV, these interactions brought

about significant advections of IPV associated with

mutual confluence of AE6-cyclones.

The computation of the fraction of IPVexplained by

the IVC� 1 showed that the latter term generally

contributed less than F 6% to the total field of IPV

(Fig. 11i and j), as expected due to the relatively large

radii of curvature and the overall low IVC� 1 values.

The largest negative dissimilarities between the two

Fig. 10. Cross sections of potential vorticity. Colorbar units are 10� 9 m� 1�s� 1: (a) zonal section along transect I for cruise leg 1 (northern half

of the domain); (b) idem for the meridional section along transect II; (c) meridional section along transect IV for cruise leg 2; (d) idem for zonal

section along transect IV. Density contours are superimposed for reference in all plots.


fields were observed at the swoddy center, where the

largest anticyclonic IV (f� 0.05) accounted for

� 6% difference between IPV and SPV (Fig. 11i and

j). Additionally, some 5% positive difference between

IPV and SPV featured in the confluence zone between

AE6 and the SC during leg 2. Only in these cases did

the relative vorticity term play a relatively significant

role in Eq. (1).

Fig. 11. Horizontal distributions at the 27.02 rh isosurface of: (a) pressure field (dbar) for leg 1, (b) layer thickness field (m), (c) isentropic

vorticity (IVC� 1), (d) Rossby-Ertel IPV field (10� 9 m� 1�s� 1). Velocity vectors at the 27.02 rh isosurface are superimposed for reference. (e–

h) idem for leg 2; (i) departure (%) of the IPV field from the planetary contribution or SVP (SPV= g� 1CN 2) for leg 1; (j) idem for leg 2.


The IPV and SPV fields appeared essentially sim-

ilar, but in areas where the NC and SC showed the

most intense interactions against the swoddy core

(along the AE6 periphery) the geostrophic flow did

not follow IPV isopleths, and seemed to advect IPV.

In these cases, the largest positive advections of IPV

by the geostrophic flow occurred at the NC and SC

locations. A fluid particle following the flow and

limited by the same density and PV values must adapt

its motion in order to conserve its PV, for which

vertical and horizontal ageostrophic motions may be

necessary, and the flow must behave quasi-geostroph-

ically in the 3D field.

During leg 1, the geostrophic flow circulated from

the zonal to the meridional transect crossing over a

high-IPV zone at the NWof the domain, caused by the

NC. Negative IPV advection from 0.11 to 0.13�10� 9

m� 1 s� 1 over f 15 km as water parcels traveling with

the geostrophic flow was observed (Fig. 11d). To

examine the 3D behavior of the flow, a pair of water

parcels delimited by the IPV lines 0.045 and 0.05�10� 9

m� 1 s� 1 and density between 27.01 and 27.03 was

colored in Fig. 12a and b. It can be observed that,

between transects I and II, the lower bound of these

water parcels was forced f 10 m upwards from 140 to

130 dbar in a presumed vortex shrinking process. In

fact, the decrease of layer thickness (or the equivalent

decrease of stratification) as water parcels traveled

isentropically with the geostrophic flow (Fig. 11c)

provided evidence that they were forced to shrink their

vortex lines as the flow around AE6 approached the

NC.

Fig. 11 (continued).


Conversely, the opposite situation was observed at

the southern AE6 front during leg 2. In this case, the

geostrophic flow seemed associated with a sequence of

positive (from 0.16 to 0.08�10� 9 m� 1 s� 1 over f 30

km, at the SE) and negative (from 0.05 to 0.08�10� 9

m� 1 s� 1 over f 25 km, at the SW) IPV advections.

The sketch presented in Fig. 12c through e depicts three

sections perpendicular to the main flow at the S AE6

boundary, at approximately 6.1 (V), 6.3 (VI) and

6.5jW (VII), respectively (see Fig. 1 for location). In

this case, the water parcels delimited by the isopycnals

27.01–27.03 and the 0.05 and 0.055�10� 9 m� 1 s� 1

IPV lines were colored. One can observe that the lower

limit of the colored box appeared to deepen f 20 m

from Fig. 10e through c (the direction of the flow is

from e towards c). It can be inferred that from the

vertical sections VandVII downwardmotionmust take

place below the 27.02 surface, in response to the

positive IPVadvection by the geostrophic flow. Hence,

as the geostrophic advected negative IPV at the SW

front, downward stretching of water parcels could be

inferred. From these results, two inferences can be

Fig. 12. Detailed cross sections of potential vorticity (10� 9 m� 1�s� 1): (a) zonal section along transect I for cruise leg 1 (northern half of the

domain); (b) idem for the meridional section along transect II; (c) meridional section along transect V for cruise leg 2 (southern half of the

domain); (d) idem along transect VI; (e) idem along transect VII. Density contours are superimposed for reference in all plots.


drawn: (i) either the front is not in equilibrium or (ii)

there are vertical velocities that displace water particles

diapycnally (Pollard and Regier, 1992).

The paper of Viudez et al. (1996) presented an IPV

balance for the Atlantic Jet in the Alboran Sea. These

authors studied a flow advecting PV=� 1.5�10� 9 to

PV=� 2.0�10� 9 over 30 km with a zonal velocity

ug =� 40 cm�s� 1, from which they inferred vertical

velocities O(10) m day� 1. The motion associated with

AE6 scaled approximately one order of magnitude

smaller than observed by Viudez et al. (1996). Hence,

it can be inferred that vertical motion will be very

weak and restricted only to subsurface layers. If local

changes of IPV are neglected the order of magnitude

of the vertical forcing can be estimated as follows.

Considering f 20 days as a reasonable rotation

period, the isentropic flow would need some 4 days

to cross from the zonal to the meridional section

presented in Fig. 10a and b. In this case, upward

vertical velocities of 2–3 m day� 1 could feature in

the AE6-NC interaction. Conversely, at the southern

swoddy boundary (Fig. 10c through e), water parcels

traveling isentropically are observed to descent ap-

proximately 20 m over 35 km, which would imply

downward velocities of f 5–7 m day� 1, associated

with AE6–SC interactions.

3.5. Vertical forcing

The theory of baroclinic waves predicts upward-

downward vertical motions as a response to changes

in flow curvature (Palmen and Newton, 1969, p. 144;

Holton, 1979, p. 136). According to the QG theory,

the vorticity field may change in response to stretch-


ing and shrinking processes by ageostrophic vertical

motions.

In the Omega equation (Holton, 1979) for strongly

meandering flows, the term that includes the advec-

tion of geostrophic vorticity (� vg�jhfg) plays the

main role in forcing vertical motion. Viudez et al.

(1996) discussed the importance of the advection of

density in the Omega equation, since both effects are

opposite and cancellation of both terms makes the

quantitative estimation of vertical motion unfeasible.

To overcome this, the Omega equation was applied as

introduced by Hoskins et al. (1978), in which this

cancellation effect is solved by the formulation of the

Q vector:

N2r2hwþ f 20 B

2w=Bz2 ¼ 2fj � Q; ð5Þ

where

Q ¼ g=R0½ðBug=BxÞðBUV=BxÞþðBvg=BxÞðBUV=ByÞ;ðBug=ByÞðBUV=BxÞþ ðBvg=ByÞðBUV=ByÞ: ð6Þ

Convergence of Q indicates regions where upward

motion is taking place (Hoskins and Pedder, 1980),

and the RHS of Eq. (5) represents the vertical velocity

forcing term.

To contrast the vertical motion inferred in Section

3.4, the divergence of Q at the 27.02 surface is

presented in Fig. 13. It was observed that the strongest

vertical activity was restricted to regions associated

with the confluence of the NC and SC with AE6. These

plots showed conspicuous upward forcing although the

larger values for | |j�Q| were associated with periph-

eral circulation around the swoddy, with j�Q>F 0.3�10� 16 m� 1�s� 3. During leg 1, the most in-

tense upward forcing was inferred at the N swoddy

boundary, where divergence of Q attained the maxi-

mum negative values of � 0.3�10� 16 m� 1�s� 3. Simi-

larly, during leg 2, the most intense downward forcing

was observed at 45.0jN 6.1jW and divergence of Q

exceeded + 0.4�10� 16 m� 1�s� 3. Both estimations are

coherent with the inferences drawn from IPV consid-

erations in the previous section.

Hence, the location of active areas in terms of

vertical forcing at a density level crossing the swoddy

core was consequently coherent with the previous IPV

analysis. They were found in relation to AE6–NC and

AE6–SC interactions and inferred meanders. These

resulted in the strongest vorticity advections in the

domain for both cruise legs. According to the com-

putation of j�Q, the term involving density advection

seemed to play a minor role and upward vertical

motion took place when positive vorticity advection

occurred (and the term � vg�jh�fg > 0) and vice versa

when there was negative vorticity advection (and the

term � vg�jh�fg < 0) (e.g., Holton, 1979).

3.6. Error estimates

The accuracy of observations is determined by

errors committed by the combined effects of both

instrumental errors and by limitation of the construc-

tion of the 3D field from scattered and non-synoptic

observations. Effects of spatio-temporal resolution can

be held responsible for errors of up to 50% on the

diagnosis of vertical velocities with the Omega equa-

tion (Allen et al., 2002). Estimates of the spatial

distribution of statistical analysis errors from the

associated OSI solution to the SC interpolation meth-

od (Franke, 1988) were computed as in Gomis et al.

(2001) (not shown). Large errors concentrated near

the boundaries and in data voids. In general the center

of the domain was dominated by errors lower than

5%. Areas with error-to-signal variances larger than

10% (for dynamic height) were blanked in the pre-

sented maps.

The error in the calculation of dynamic height at

rh = 27.02 can be estimated as e/c 0.02 dyn cm (cf.

Fig. 9b, c and d). Derived fields have been calculated

at each grid point using central differences from the

dynamic topography. The error in the geostrophic

velocity (using a scheme of finite differences over a

grid arm c 7.9� 11.1 km2) can be calculated:

eU = e//CDxc 2�10� 3 m�s� 1. Similarly, the error in

the computation of relative vorticity from a second

derivative of the geopotential field can be calculated

as (Atkinson, 1989, p. 319) e1/C= 4e//C2Dx2 + 4e//

C2Dy2c 0.02. The total error-to-signal variance is of

the order of 30–60% for the IV (see Fig. 12). The

error in the buoyancy term of the IPV equation�[� q� 1�Dq/Dp] can be estimated. Using a window

size = 20 m, ebuoy =� q� 1�eq/Dzc 10� 6 m� 1, if we

consider the instrumental error in the determination of

the density eqc 0.02 kg m� 3. Hence, the error

estimated for the IPV can be approximated by

Fig. 13. Vertical forcing as estimated from the divergence of Q vector (10� 16 m� 1�s� 3) at 27.02 rh; shaded areas and dotted contours for

divergence (downward forcing) and lines for convergence (upward forcing): (a) for leg 1, (b) idem for leg 2. Dynamic topography at 27.02

rh (27.17 rh reference level) contours are superimposed for reference. The colorbar represents positive values or divergence of Q. Unit is

10� 16 m� 1�s� 3.


eIPV={[ebuoy�IV]2+[(q� 1�Dq/Dp)eIV]2}1/2 < 10� 11 for

IV values of 0.04C and 4�10� 6 m� 1 for the buoyancy

term.

Additionally, the error in the jQ term may be

assessed if we consider that the error in the term UyyVz

is the dominant contribution to its computation (Viu-

dez et al., 1996). e(Vz) can be obtained from the

thermal wind relation (�V,U)z =j( g/Cq): eVz = geq/

CqDxcO (10� 4) s� 1. Similarly, eUyy = 4�eU/Dy2cO(10� 10) m� 1�s� 1. Then, eUyyVz=[(eUyy�Vz)

2+

(Uyy�eVz)2]1/2cO(10� 14) m� 1�s� 2 for values of Uyy

(10� 10 m� 1�s� 1) and Vz (10� 4 s� 1) giving 2CeUyyVzc 0.02�10� 16 m� 1�s� 3, which implies a

notable error in the forcing term of the Omega

equation.

Thus, the determination of the vertical circulation

is be contaminated by propagating errors in the order

of 25–75% (e.g., Viudez et al., 1996; Allen and

Smeed, 1996; Gomis et al., 2001). Estimation of

errors of analysis for dynamic height, isentropic

vorticity and isentropic potential vorticity showed that

these fields (computed from CTD data) should allow

the diagnosis of both the sign and order of magnitude

of the features. However, the vertical forcing in the

Omega equation shows larger errors and it might be

difficult to use the results for quantitative verification,


as Gomis et al. (2001) observed in the specific context

of the Alboran Sea.

4. Summary and conclusions

A mesoscale swoddy (AE6) that had presumably

detached from the IPC before summer 1998 was

sampled with a CTD grid in the Bay of Biscay in

August 1998. Information from T/P-ERS SLA altim-

etry measurements showed evidence of an anticy-

clonic anomaly at the central Bay of Biscay from

July until September 1998. Thermal structure and

altimetry data both agreed in explaining the features

within the sampling domain. Linked to the AE6, a

pair of cyclones appeared attached at its N and S

boundaries (NC and SC). A sequence of SLA maps

showed the whole tripolar system subjected to a SW

drift at a rate of 1.5 km per day, variable between 1.0

and 3.5 km/day. This fact appeared to be supported

by a maximum thermohaline drop at 400 dbar at

rates of 0.045 and 0.005 jC day� 1 at the swoddy

center.

Outcropping of the thermocline characterized the

central region above the swoddy core and spiciness

values significantly lower than in the undisturbed

surroundings were observed. Onken’s (1990) pro-

posed mechanism of creation of reversed baroclinicity

was considered for the explanation of the doming of

the isopycnals below the seasonal thermocline. The

swoddy core featured a vertical salinity profile with

constant 35.70 salinity between 80 and 200 dbar.

Potential temperature decreased slightly with depth,

yielding a weak gradient from 12.75 to 12.55 jC. Atthe level of the AE6 core (f 130 dbar), the swoddy

was well defined by higher spiciness values than in

the surroundings, hence being evident as a warmer,

saltier and lighter homogeneous nucleus with dimen-

sions f 50 km diameter.

As well as the effect of steering currents, two

other disturbing agents were found. Firstly, the

swoddy was disturbed by strong short period internal

waves that introduced a strong range of variation

over the upper layers. Thus, the uppermost layer

could not be used for dynamic considerations. Sec-

ondly, interactions against mesoscale cyclones played

a determinant role in defining the vorticity field of

the geostrophic flow. The core boundary appeared

intensified by interactions against the aforementioned

NC and SC.

The dynamic picture showed a central anticyclonic

structure surrounded by a number of conspicuous

structures making up the swoddy system in coherence

with the satellite data. Dynamic fronts appeared in-

tensified and deflected as a response to the interfer-

ence swoddy against NC and SC. As a result, the

largest advections of geostrophic vorticity were in-

ferred associated with the flow around this tripole

structure. Mean velocities of O(15) cm s� 1 were

estimated as characteristic of the flow around AE6,

whereby Rof 0.01 and rotation period of 15–20 days

was inferred. The flow around the swoddy from the

subsurface to intermediate levels of O(500) dbar was

found to be approximately in hydrostatic and geo-

strophic equilibrium.

Geostrophic relative vorticity values were general-

ly negligible with respect to unity, and the IPV field

resembled the stratification field. Geostrophic dynam-

ics may satisfy the requirement of PV conservation

and represent the 2D real flow. However, localized

vertical forcing was inferred at locations affected by

AE6 interactions against the surrounding cyclones. At

these locations, the geostrophic flow followed neither

stratification nor IPV isopleths. It seemed to advect

PV isentropically, for which vertical and horizontal

ageostrophic motions may be necessary and small

departures from geostrophy were inferred. The IPV

analysis performed at the 27.02 isopycnic surface

revealed that the IPV showed little departures with

respect to the corresponding planetary contribution or

SPV, except at the AE6 core and at the S boundary.

However, due to the weakness of the velocity field

these deviations were smaller than cited in the litera-

ture. Vertical motion was therefore inferred to be

relatively weak, between 2 and 7 m day� 1, restricted

only to the periphery and more intense at the SW

front. The vertical forcing at 27.02 isopycnic surface

as calculated from the Omega equation was coherent

with the IPV study.

Thus, the swoddy may be understood as a system

with an upper, non-linear layer affected by internal

wave turbulence. Exchange of heat and momentum

through the boundary layer seemed to provide energy-

sinking mechanisms along boundaries and tempera-

ture losses by turbulence diffusion, all of which are

enhanced by incipient cooling and mixing. Further, at


the level of the AE6 core ageostrophic forcing by AE6

interactions against the surrounding structures are

proposed as another source of instability, triggering

ageostrophic mass and momentum exchanges be-

tween the swoddy and peripheral eddies. On the

frontal scales of mesoscale eddies, the conservation

of potential vorticity implies vertical motions when-

ever the vorticity of the fluid changes. While fronts

hinder horizontal transfers of heat, momentum and

other properties, they play a crucial role in enhancing

vertical exchanges. External interactions of eddies

with other structures are the source of three-dimen-

sional vorticity changes associated with the velocity

field and are therefore relevant in considering the

conservation of IPV.

From the IPV analysis, it was inferred that only in

regions with interactions between structures did

relative vorticity play a significant role. Even in

these cases, absolute departures of IPV from its

corresponding SPV were always less than 10%.

However, it was only in these areas where vertical

motions were inferred from QG dynamics consider-

ations. Elsewhere over subsurface layers, a rather

geostrophic, linear and stationary gyre prevailed. It

can be concluded that the stability of the swoddy

system at the level of the core appeared to be

strongly dependent on the degree of interaction

between the central anticyclone and the satellite

cyclones.

Acknowledgements

This work was supported by contract CC-96-

MAR-1872-C0301 from CYTMAR and Instituto

Espanol de Oceanografıa. Cooperating institutions

(University of Vigo and University of Oviedo) are

also acknowledged. The authors would like to

express their gratitude to the crew of R/V Professor

Shtokman and M. Blanco who helped in data

acquisition. The Deutsches Zentrum fur Luft-und

Raumfahrt (DLR) processed NOAA AVHRR data

and were obtained through the public access gateway

(http://isis.dlr.de/). SLA maps were provided by CLS

though the AVISO web gateway (http://www-avi-

so.cnes.fr). We kindly acknowledge Dr. L. Valdes and

the project ‘Studies on time series of oceanographic

data’, managed by the IEO. One of the authors (RS)

was supported by IEO and from European Union

FSE funds, and is currently supported by ATOMS

Project (FCT contract PDCTM/P/MAR/15296/1999).

Comments by three anonymous referees were helpful

in improving an earlier version of the manuscript. We

are indebted to Mr. B. Morris for his valuable

suggestions.

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