Download Year-long float trajectories in the Labrador Sea Water of the eastern

Deep-Sea Research II 46 (1999) 165 — 179
Year-long float trajectories in the Labrador Sea
Water of the eastern North Atlantic Ocean
Kevin G. Speer *, John Gould, Joseph LaCasce
Laboratoire de Physique des Océans, IFREMER/CNRS, B.P. 70, 29280 Plouzané, France
Southampton Oceanography Centre, Southampton, UK
Received 12 December 1997; received in revised form 1 June 1998
Abstract
Year-long Lagrangian trajectories within the Labrador Sea Water of the eastern North
Atlantic Ocean are analysed for basic flow statistics. Root-mean-square velocities at 1750 m
depth are about 2 cm/s, except within the North Atlantic Current, where they are twice as large.
These values are consistent with previous Eulerian measurements and extend those results to
a much larger domain of the eastern basin. Mean flow estimates in boxes large enough to
contain about 1 float-year of data indicate that Labrador Sea Water, having crossed the MidAtlantic Ridge (not resolved) near 50—55°N, presumably with the North Atlantic Current,
partially recirculates to the north in the subpolar gyre, as well as entering the subtropical gyre
and continuing south and west. The circulation of this water mass, as defined by the 1 yr
average velocities, is stronger than traditional models of deep circulation would suggest, with an
interior flow of roughly 1 cm/s. Mean speeds up to 3 cm/s were observed, with the highest
values near the Azores Plateau. North of 45°N—55°N, mean eastward speeds closer to 0.2 cm/s
were observed. Wind-generated barotropic fluctuations may be responsible for some part of the
transport at this depth. 1999 Elsevier Science Ltd. All rights reserved.
1. Introduction
The spreading of water masses in boundary currents appears to be the main
mechanism for the transport of anomalous properties far from their source in the
ocean. The flow of different water masses along the western boundary of the Atlantic
Ocean, bringing Southern Ocean water masses northward and North Atlantic water
masses southward, is a well-known example of this transport (Warren, 1981). What
are less well known but of critical importance to theories of the evolution and
*Corresponding author. Fax: 0033 2 98 22 44 96; e-mail: [email protected].
0967-0645/99/$ — see front matter 1999 Elsevier Science Ltd. All rights reserved.
PII: S 0 9 6 7 - 0 6 4 5 ( 9 8 ) 0 0 1 0 3 - 9
166
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
maintenance of stratification and property distributions are the flow and mixing of
water masses in the interior of the ocean.
A number of water masses are formed in the North Atlantic Ocean, owing to
intense air—sea interaction in polar and subpolar regions, and to the salinification
occuring in the Mediterranean basins. The generation process and ultimate temperature, salinity, and depth of each water mass in the interior of the ocean are coupled
together and depend on the nature of the mixing processes governing their evolution.
Labrador Sea Water formation occurs at the dense end of the range of subpolar water
masses, and is therefore one of the principal mechanisms for removing heat and salt
anomalies from further interaction with the atmosphere. The spreading of this water
mass in the interior has received much attention, with observations showing an
influence on the large-scale distribution of potential vorticity (Talley and McCartney,
1982), and time variations in property transports (Sy et al., 1997) thought to be related
to atmospheric variability (Dickson et al., 1996). The persistence of a water mass (or
property anomalies that distinguish one water mass from another) over decades in the
presence of variability in its source leads to the question of a dissipation or water mass
consumption time scale. To begin to address this issue, the basic space and time scales
of circulation need to be measured.
The flow of Labrador Sea Water was estimated by Saunders (1982) to be to the
south and west in the eastern basin from 32 to 55°N, with speed decreasing from
a maximum near 1 cm/s (14;10 m/s net across the eastern basin) in the north to
0.1 cm/s (2;10 m/s net) in the south. His results were based on transport calculations using hydrographic sections across the eastern basin, adjusted so that the
vertically integrated transport agreed with the local wind-stress curl. Direct measurements presented here show a qualitative agreement in the southern part of the basin
with Saunders (1982), though of very different strength, and a deep extension of the
North Atlantic Current in the north not present in his result. Cunningham and Haine
(1995a) describe a circulation qualitatively similar to ours, from widely-spaced hydrographic stations in the eastern North Atlantic. They emphasize the time-dependent
nature of the supply of Labrador Sea Water to the region, and the consequences this
has on the interpretation of tracer distributions. A range of water mass ages from
5—20 yr is estimated, based on time series of salinity on an isopycnal surface.
Paillet et al. (1997) have described the flow of this water mass in some detail, using
an inverse model to incorporate hydrographic and wind data on a uniform grid. They
infer a transport of about 6;10 m s entering the eastern basin, with about one-half
of this circulating to the south and west within the eastern basin. Low salinity
suggestive of a narrow, more intense flow around the Azores Plateau is described as
well. Our estimate of southward transport from the zonally averaged meridional float
velocity at 45—50°N (!0.5$0.3 cm/s ) is 2.8;10 m s for a 500 m thick layer
(Section 4), which agrees with their value within error.
With the questions of dispersion and water mass consumption in mind, the
Eurofloat experiment was conceived to study the flow and mixing of Mediterranean
Water and Labrador Sea Water in the eastern North Atlantic Ocean. Only results
from that part of the experiment dedicated to the Labrador Sea Water are used here;
the experiment is ongoing, and the purpose of this note is to discuss the space and time
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
167
scales of motion from year-long Lagrangian trajectories, following this water mass
after it has crossed the Mid-Atlantic Ridge and entered the eastern basin.
The direct measurements reported here show a flow structure consistent with both
the recirculation of Labrador Sea Water in the subpolar gyre and its incursion into the
subtropical gyre, and further southward movement. The year-long mean strength of
this circulation is observed to be significant over more than half of the domain, where
mean speeds of up to several centimeters per second are found.
2. Data
Twenty-one MARVOR subsurface floats were deployed, nominally at 1750 m
depth, in September 1996 from the B.O. D’ENTRECASTEAUX. The floats were
programmed to maintain actively their depth at the 1750 db isobar (close to 1730 m)
to within $50 db. They return to the surface every three months to transmit the data
via the ARGOS network. The design and operating characteristics of the MARVOR
float are described in Ollitrault et al. (1994). Out of the twenty one floats, one never
transmitted any internally recorded data, and only three-month displacements are
available from the ARGOS position at the surface; two other floats ran aground on
the Rockall Bank and only partial trajectories are available. Altogether 18.6 floatyears of data were obtained from the first year.
The data consist of an acoustic time-of-arrival signal used for tracking, as well as
temperature and pressure. The acoustic information is recorded once per day, while
the temperature and pressure values are both an average of 24 measurements sampled
at one per hour from the preceeding day. Tracking accuracy depends on many
parameters, including the distance and bearing to the sound sources, speed of sound
variability, clock drifts, and data return through the ARGOS network. These and
other technical issues are described completely in Menzel (1998). The final root-meansquare position accuracy is thought to be $2 km. Temperature and pressure are
accurate to $0.01°C and $20 db (with precision $1 db).
A sound source network was installed and maintained by the Service Hydrographique et Oceanographique de la Marine; tracking also depends on further sound
sources deployed in collaboration with the Institut für Meereskunde Kiel, the University of Rhode Island, and the Woods Hole Oceanographic Institution.
The year-long trajectories are broken into four three-month cycles, separated by
a 47$2 h transmitting period at the surface, during which time the floats can drift
a substantial distance away from their initial position. Depending on the quantity to
be derived from the track, this gap may be ignored or taken into account. The
deployment was made on a regular grid of about 5° longitude by 2° latitude (Fig. 1),
except at the northern edge of the domain, where deployments were made more-orless evenly spaced along the ship track southwest of the Rockall Bank.
Breton for ‘‘SEA-HORSE’’.
168
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
Fig. 1. Year-long trajectories from MARVOR floats deployed in the Labrador Sea Water at 1750 m depth.
The trajectories are broken into four cycles of 3 months each.
3. Flow statistics
The strongest persistent flow occurs in an anticyclonic eddy at 21°W, 52.5°N, with
a maximum speed of about 15 cm/s, and typical speed of 5—10 cm/s. Based on closed
loops in the trajectory, the float appears to have been trapped for only two revolutions, of period 13 days and diameter 50 km, but velocity shows it to have been
within the eddy flow field for the entire 90 day cycle, oscillating at the 13 day period.
No other float in the area showed similar behaviour, and this single trajectory has
a strong effect on statistics at its latitude. Overall, 5% of the daily speed records are
greater than 10 cm/s, up to a maximum of 20 cm/s, and about 30% are greater than
5 cm/s. The median speed is 3 cm/s. This value is close to the rms velocity
u"v"3.2 cm/s. Mean velocity of the full data set (all floats) is weaker,
U"(!0.09, !0.37) cm/s.
These very simple statistics demonstrate immediately that the flow at mid-depth is
moderately energetic over a broad geographical range. Local measurements from
moorings in the NEADS 5 (46°N, 17°W; Gould, 1983), and TOURBILLON (47°N,
15°W; Mercier and Colin de Verdiere, 1985) experiments both show the same rms
velocity within error near 1500 m depth. A more complete description of the eddy
energy distribution along 48°N (Colin de Verdiere et al., 1989) shows similar energy
levels at this depth. Their Eulerian integral velocity time scales at 48°N (evaluated at
the first zero crossing of the autocorrelation) were all close to 15 days at 1500 m,
except at the 25°W mooring cluster, where it was 25 days. Variance or rms velocity
was the same as the Eurofloat value within error up to 30°W, before increasing further
west as it entered the North Atlantic Current. Variance increases to the north
(Schmitz and Hogg, 1978) also because of the presence of the North Atlantic Current.
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
169
Table 1
Flow statistics as a function of latitude. Velocity units are cm/s and Lagrangian time scales ¹x and ¹y in
days
°Lat.
float-yr
º
»
º-error
»-error
rms u
rms v
¹x
¹y
53—56
50—53
47—50
44—47
41—44
38—41
35—38
3.98
2.09
3.38
4.09
3.05
0.98
0.59
0.22
0.10
0.22
0.22
!1.19
0.69
!2.24
0.40
0.09
!0.45
!0.59
!0.57
!2.33
!0.80
0.29
0.51
0.36
0.34
0.51
0.42
0.91
0.41
0.42
0.37
0.35
0.43
0.56
0.51
4.53
3.84
2.60
2.18
2.04
2.75
2.43
4.27
4.20
2.91
2.24
2.26
2.81
2.68
7.3
7.8
8.4
7.5
10.0
7.6
14.0
9.1
7.8
8.4
8.2
8.0
8.4
6.8
Nevertheless, a possible connection between increasing energy and increasing wind
stress to the north was suggested by Mercier and Colin de Verdiere (1985).
Both the TOPOGULF and TOURBILLON experiments found a substantial
vertical coherence of velocity, and a rough equipartition of energy between the
barotropic mode and the first baroclinic mode. The long time series of the IFM Kiel
mooring 276 near 33°N also shows a similar mix of barotropic mode (60%) and first
baroclinic mode (Müller and Siedler, 1992). The first mode zero crossing occurs near
1300 m depth over much of the eastern basin at Lats. 40—50°N.
A mid-depth intensified empirical mode in the TOPOGULF data captured what
was thought to be the cause of the greatest variability near 1500 m depth in the
eastern basin in the form of strong cyclonic events. These high-potential-vorticity
eddies were thought to originate in either the Charlie—Gibbs Fracture Zone and
North Atlantic Current to the north, or in the Mediterranean Water to the south;
however, they were seen preferentially west of the ridge, and closer to the North
Atlantic Current.
The basic statistics are presented as a function of latitude (Table 1) since the
Eurofloat dataset is still rather sparse, and the meridional extent of the domain is
nearly twice the zonal extent. Such a presentation may distinguish subtropical and
subpolar regimes in the eastern basin. Means (º, ») are larger in the south, and
although the amount of data is less, relative errors are smaller. A fairly consistent
pattern of northward and eastward flow at latitudes north of 50°N, and southwestward flow in the south appears. Mean flows reported by Mercier and Colin de
Verdiere (1985) were, in contrast, to the northwest, but they noted a large degree of
inhomogeneity in the TOURBILLON cluster. This is an arguement for Lagrangian
measurements, which average over spatial variability (though our year-long trajectories are not long enough for confidence at many locations).
Variance is greatest in the northern boxes, with u:v:4 cm/s, close to or within
the North Atlantic Current, and quite uniform at other latitudes, with
u:v:2—3 cm/s. The Lagrangian time-scales (¹x, ¹y) show little spatial variation
and are 7—10 days, except for the southwestern corner of the domain, where the flow
around the Azores Plateau shows scales of (14, 7) days.
170
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
Table 2
Diffusivity (m/s) and derived time (days) and space (km) scales. Errors in diffusivity are approximately
$40% at 20 day lag
°Lat.
K
V
K
W
¹)
V
¹)
W
¸)
V
¸)
W
53—56
50—53
47—50
44—47
41—44
38—41
35—38
510
870
673
736
1243
275
775
1043
572
745
801
870
489
243
5.8
13.7
23.1
35.8
69.1
8.4
30.4
11.8
9.0
25.6
38.9
48.4
15.0
9.5
22.5
45.4
51.8
67.4
121.9
20.0
63.8
46.0
29.8
57.4
73.4
85.3
35.7
20.0
Fig. 2. Dispersion (km) versus time for the Eurofloat data (left) and the Eastern Basin experiment (right;
Spall et al., 1993).
Diffusivity values (Table 2), calculated by integrating the velocity autocorrelation
and taking the result at 20 day lag, differ by a factor of 4—5 over the region, but
without a simple structure. Time and length scales based on diffusivity and variance,
¹)"
2K 2K
V,
W
u v
¸)"
2K 2K
V
W
u v
are given, and suggest longer time and space scales in the 41—44°N band. This band is
a region of persistent southwest flow in virtually all tracks.
Float absolute dispersion (Fig. 2) was calculated for the Eurofloat data, and for data
from the Eastern Basin experiment (Spall et al., 1993), defined as
D "1(x!x !1x!x 2)2
V
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
171
where x is the initial longitudinal position, brackets mean an average over all floats,
and the average position of the floats in each individual dataset has been subtracted to
define the displacements (similarly for D ). Floats in the Eastern Basin experiment
W
were launched in the region 20—35°N, to the south of the Eurofloat experiment, at
shallower depths in the water column (nominally 1100 m). Floats deployed in
Meddies have been excluded from this calculation, and for both data sets, no fewer
than 8 floats were used for averages at later times.
The Eastern Basin floats show a marked zonal anisotropy in their spread, as
discussed by Spall et al. (1993); in contrast, the Eurofloat dispersion is nearly isotropic,
suggesting that the meridional spreading is as great as zonal spreading for these floats.
In addition, the Eurofloat dispersion increases roughly linearly in time after approximately 30 days, a signature of a diffusive process under certain conditions (Taylor,
1921). The rate of increase in the zonal dispersion of the Eastern Basin floats is greater,
growing approximately as t. Moreover, there is a suggestion of a cessation of
spreading in the meridional direction. This apparent saturation occurs late, when
there are fewer floats, and only after one year, which is the duration of the present
Eurofloat dataset, so we cannot compare this aspect.
In any event, it is clear that the statistics for the two datasets are quite different,
despite the fact that they are both at mid-latitudes of the eastern basin, below the
thermocline. We do not think the differences in depth or latitude explain the different
dispersion relations, but rather that dispersion in both regions is controlled to some
extent by topography as well as the b-effect, and that a measure of dispersion that
takes this into account shows similar behaviour (LaCasce and Speer, 1998; LaCasce
et al., 1998).
4. Deep circulation
The domain was divided into 5°;3° boxes to construct average quantities and
averaged time-series (Fig. 3). There are 23 non-empty boxes, but only 23 have more
than a cut-off of 30 float-days of data. The average number of data in a box is
310$180 float-days, while the range is 50—628 float-days. Estimates of error (Davis,
1991) are included with the mean in Fig. 3; a principal value decomposition was not
significantly different in appearance and is not shown. The overall structure suggested
by the zonal box-averages (Table 1) stands out. Evidently the Labrador Sea Water
crosses the Mid-Atlantic Ridge (though this inflow is not resolved) at 50—55°N,
presumably with the North Atlantic Current, part of it recirculates to the north in
the subpolar gyre and part of it enters the subtropical gyre and continues south
and west. A simplified, zonally averaged mean flow (Fig. 3b) illustrates the overall
structure.
The inflow was clearly present in the inverse model of Paillet et al. (1997) at
latitudes of 51—52.5°N on the Mid-Atlantic Ridge. We placed one float in the
Charlie—Gibbs Fracture Zone at 52.5°N to try to sample this inflow, but unfortunately
only the net three-month displacements are available from this float, and these have
been interpolated and included in the box-mean velocity. The mean velocity from this
172
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
Fig. 3a. Mean velocity and error estimate. b. Mean velocity from a low-order polynomial fit to illustrate
the large-scale structure as a function of latitude.
float alone for the first three months is (u, v)"(2, 0) cm/s. At later times it turns north
and moves at an average speed of 3—4 cm/s.
Southwestward flow in the subtropical gyre would be of little intrinsic interest if the
floats were within the thermocline. At 1750 m they are well below it, at a level where
traditional models based on the idea of a deep upwelling into the thermocline
(Stommel and Arons, 1960) would predict northeastward flow. A modified deep
circulation scheme is used below, thought to be more appropriate to the mid-depth
layer of the North Atlantic Ocean.
Further measurement ought to show whether or not the observed flow around the
Azores Plateau (35—40°N) continues west as a zonal jet or follows the topography
back to the north. That these few data near the Azores Plateau actually do represent
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
173
a significant southward and westward flow of deep water is supported by the
observation of low salinity next to the plateau relative to the interior (Wilson
et al., 1995; Paillet et al., 1997) and a net westward geostrophic flow on the southern
flank of the plateau, above the crest of the ridge at 36.5—37°N (Wilson et al.,
1995).
Data coverage across the basin is best at 45—50°N, and the estimate of zonally
averaged meridional velocity in this latitude range, 0.5 cm/s (Table 1), may be used to
find a value for the meridional transport. Assuming that this number applies to a layer
500 m thick (or is the maximum value of flow in a somewhat thicker layer) across the
entire basin at 48°N (1137 km), the southward transport into the eastern basin at this
latitude is 2.8;10 m/s.
4.1. A modified Stommel—Arons flow
Buoyancy forcing forces directly the mid-depth circulation, and this is traditionally
represented by the Stommel and Arons (1960) model. A variation of this model is
displayed here (Fig. 4) to illustrate the circulation that would be expected in this
simple framework. The model is of a single homogenous mid-depth layer,
1000—3000 m depth, with realistic topography and frictional boundary layers to close
the circulation in the horizontal plane (Speer et al., 1993). The forcing is a uniform
downwelling, of magnitude w"!2.5;10\ cm/s, the opposite of the usual model,
to reflect the mid-depth compression and southwest interior flow expected in a basin
with entraining overflows (e.g. the Denmark Strait overflow) from marginal seas,
which cause the basin-averaged vertical velocity to decrease with height at mid-depth
Fig. 4. Deep circulation forced by uniform cross-isopycnal velocity (compression) in a homogenous layer
below 1000 m depth. A sink of mass (5;10 m/s ) is prescribed in the southwest corner equal to the input
through the upper interface. Velocity scale as in Fig. 3a.
174
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
(Speer, 1993). This decrease simply reflects the mass lost to the layer by entrainment
into the boundary current, and for this value of w amounts to about 5;10 m/s.
A dominating aspect of the model’s interior circulation is the northward flow along
the western flank of the Mid-Atlantic Ridge, but there are no direct measurements
near the Azores Plateau to support it. The primary boundary current at the western
boundary shows increasing transport to the south as it is fed mass by the interior.
A recirculation occurs in the northeast sector, north of 45°N, in which water moves
south in the interior and north along the eastern boundary. A single float trajectory
near the eastern boundary at 52°N shows southward flow (Fig. 3a), not northward
flow, but the eastern boundary is scarcely sampled by the year-long dataset.
Even with what seems like a plausible forcing, the interior model flow is weak, of
order 0.1 cm/s over the flank of the Mid-Atlantic Ridge. There is some intensification
of flow east of the Azores Plateau, as observed, but with a maximum velocity that is
much weaker than observed, by a factor of ten or so.
Strong bottom friction in the model may be the reason for the weak velocities. The
value used here, 0.01 m/s, is several times larger than that used by Greatbatch and
Goulding (1989) in their barotropic model of the Atlantic basin, and they obtained
velocities of 0.2—0.5 cm/s in the eastern basin. However, their forcing was the monthly
climatological wind stress anomaly, which led to vertical velocities several times
stronger than that used here. Even if friction could be reduced in our model, it seems
unlikely to provide for flows ten times stronger, since a simple scaling of the largescale vorticity equation (notation is standard) bv"fw for a water column 2 km thick
X
(over the flank of the ridge), b"10\ cm\/s, f"10\/s, and the above value for
w implies a v&0.1 cm/s. Amplification by flow over a sloping bottom could increase
this scale somewhat, but hardly by a factor of ten over the ridge as a whole. Finally,
away from boundaries (where depth is small) the contribution of the frictional term is
weak, about 1%.
It is possible that with more data and a longer time series the observed mean
velocities will be lower and closer to the model results (but also less statistically
significant). The nine-year measurements of Müller and Siedler (1992) at one location
show no significant deep mean flow over the entire time series, yet year-to-year means
can be strong. Nevertheless, our float measurements sample nearly an entire basin,
avoiding some problems with point measurements, and do show a relatively strong
flow for the first year. We conclude that some other transport mechanism may be
operating as well and examine alternatives next.
4.2. Alternatives
Below the direct forcing by the wind, the mid-depth ocean is nonetheless subject to
a dynamical control by the wind through the effect it has on the shape of the
thermocline. Dynamical characteristics of the large-scale equations of motion are
guided by thermocline structure, but as shown by Rhines and Schopp (1991), flow is
only present in the northeast part of an ocean basin because of forcing by strong eddy
motion somewhere (e.g. unstable frontal zones) along characteristics that penetrate
this region. Buoyancy forcing, which does force directly the mid-depth circulation, is
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
175
weak, so it would appear that an appeal to distant eddy forcing has to be made in
order to explain the strength of the large-scale mean and fluctuating flow in the
eastern basin. However, an alternative suggested by studies of the effect of wind
variability on flow (Veronis and Stommel, 1956; Willebrand et al., 1980) is that
barotropic wind-generated motion provides a background rms energy level at middepths of order 1—3 cm/s. Willebrand et al., 1980) included a Mid-Atlantic Ridge in
their model and emphasized the resulting spatial inhomogeneity of the response to
large-scale atmospheric forcing. An enhanced rms kinetic energy on either side of the
Azores Plateau was one example of such variation, presumably due to resonantly
forced topographic waves.
Evidence for ubiquitous barotropic motion from current meter arrays in the eastern
basin was noted in the previous section. The source of barotropic motion is not
necessarily forcing by the wind. For instance, Price and Rossby (1982) have described
float motion at 1300 m depth in the western basin in terms of a barotropic wave
modified by the local topography. Numerical studies suggested that the source of
this wave was likely to be the Gulf Stream to the north rather than the wind.
The general tendency of the ocean to convert baroclinic motion to barotropic
motion was invoked as well. In the eastern basin, Spall (1994) has argued that the
instability of the large-scale meridional flow may be responsible for some of the eddy
motion and subsequent eddy tracer tranport. Yet the wind alone is evidently capable
of generating realistic deep rms velocity, and for the relatively quiet parts of the
eastern basin, south of the energetic North Atlantic Current, wind is an unambiguous
source.
It remains to show that wind variability is capable of organizing a flow consistent
with our observations. To attempt to do so, flow in a homogenous ocean with realistic
topography was computed from the Isemer and Hasse (1987) monthly climatological
wind dataset. The amplitude of the motion depends in an unsatisfying manner on the
value chosen for the bottom friction in the model. Flow strength in the monthly fields
as high as 0.5 cm/s above the eastern flank of the Mid-Atlantic Ridge could be
obtained with a value (0.25 cm/s) similar to that used by Greatbatch and Goulding
(1989). At lower values, the flow was stronger, but its structure appeared noisy and
boundary currents were poorly resolved .
The difference between the flow fields for the months of September and January
provides an estimate of the advection over that period (Fig. 5). Considering only
the region of the experiment, there is an intensification of barotropic motion to
the east and northeast of the Azores Plateau, as well as over the northern part of
the Mid- Atlantic Ridge, extending south into the eastern basin. Especially east and
north of the Azores Plateau, this structure is at least suggestive of a partial control on
the trajectories by barotropic motion, perhaps generated by the wind. A demonstration that the velocity fluctuations in the float data were oriented parallel to
bathymetry or f/H, where H is the full water depth and f is the Coriolis parameter,
would be more convincing, but, as mentioned above, the orientation of the principal
axes is not much different from the error ellipses displayed with the regional mean
flow calculation, and their relation to bathymetry is not obvious with this analysis of
our dataset.
176
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
Fig. 5. Geostrophic response of a homogenous layer to wind forcing shown as the difference between
September and January flow fields. Wind data are from the Isemer and Hasse (1987) climatological atlas.
5. Discussion
Our measurements do not yet resolve the inflow of Labrador Sea Water to the
eastern basin, a possible candidate for an internal jet, but they do indicate an active
circulation in this water mass at the scale of the entire eastern basin — and as opposed
to an active flow confined to a boundary current, with a weak interior circulation. The
flow toward and around the Azores Plateau is likely to be more a result of the
convergence of the interior flow on large-scale characteristics or geostrophic contours
related to topography and stratification, than to be a traditional western boundary
current for the eastern basin, passively fulfilling the role of a mass balance mechanism
for the forced interior circulation. In dynamical terms, the flow would be largely along
geostrophic or mean potential vorticity contours rather than across. The relation
between trajectories and the large-scale potential vorticity field is supportive
(O’Dwyer et al., 1998), but there are not enough data from the area of the plateau to be
confident of this distinction.
Theoretical studies of mid-depth flow show how it can be driven by eddy fluxes,
atmospheric, and buoyancy forcing. The strength of eddy-induced flow is unknown,
but buoyancy forcing alone seems to be too weak to explain all of the observed mean
flow. Atmospheric forcing of barotropic motion does appear to be able to generate
realistic eddy kinetic energies at depth, and further study is needed with models with
higher space and time resolution to attempt to reproduce the overall structure of the
observed circulation.
It is well known that the non-linear relation between Eulerian velocity and particle
displacements can lead to significant mean Lagrangian drift even though the Eulerian
flow may have zero mean. Without having distinguished Lagrangian drift from mean
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
177
Eulerian flow in the data, nor being confident about the definition of the mean flow
field throughout the basin from only one year of data, we can do little more than make
the suggestion that wind-generated barotropic flow is important for deep circulation.
Baroclinic modes also are affected by topography, and their influence may be combined with that of the barotropic mode to confound a simple interpretation in terms of
the barotropic mode alone. The vertical structure of lagrangian dispersion should be
a key test of dynamical models.
A diffusive time scale using i:1000 m/s (Table 1) and the basin width as a length
scale is 30 yr. Similarly, an advective time scale would be 5 yr, and a scale for the
Peclet number is 6. Rather than being a stagnant dead-end for the Labrador Sea
Water, the eastern basin appears to be resupplied or ventilated in something like 5 yrs,
not unlike water masses at much shallower levels. Moreover, the interior flow in the
southern part of the region tends toward the southwest, apparently guided by an
eastward extension of the flank Mid-Atlantic Ridge near 43°N, and thus is oriented
along rather than across the large-scale salinity field (Worthington and Wright, 1970).
The implication is again that advection dominates diffusion, and unlike the Mediterranean Water at shallower depths, which is thought to be dispersed to a great extent
by eddy fluxes, the Labrador Sea Water more nearly conserves its properties as it
flows south in the eastern basin.
Clearly, Labrador Sea Water does evolve to some extent in the eastern basin,
gaining salinity perhaps by lateral mixing (Colin de Verdiere et al., 1986; Cunningham
and Haine, 1995b) or by vertical mixing (Paillet et al., 1997), but the dominant mixing
effect may occur due to the presence of the Azores Plateau, around which relatively
fresh Labrador Sea Water wraps itself and stretches into small-scale filaments.
Vertical mixing has been found to be enhanced above the Mid-Atlantic Ridge near
20°S (Polzin et al., 1997), and, if these results can be extrapolated to the North Atlantic
Ocean, the combination of both effects suggests that Labrador Sea Water consumption or dissipation could occur principally on the Mid-Atlantic Ridge.
Acknowledgements
We wish to thank N. Cortes for help in testing the floats, O. Peden, Y. Auffret, and
M. Menzel for float preparation and deployment, and the captain and crew of the B.O.
D’ENTRECASTEAUX. All float tracking was done by M. Menzel, supported by
MAST EUROFLOAT project PL920057. The sound source network maintained by
the Service Hydrographique de la Marine in the context of the ARCANE project, and
those put in place by IFM Kiel, are crucial to float tracking. J.L. was supported by
MAST EUROFLOAT project PL920057; support for K.G.S. comes from the CNRS.
References
Davis, R., 1991. Observing the general circulation with floats. Deep-Sea Research 38 (Suppl.), S531—S571.
de Szoeke, R.A., 1985. Wind-driven mid-ocean baroclinic gyres over topography; extension of the Sverdrup
relation. Journal of Marine Research 43, 793—824.
178
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
Colin de Verdiere, A., Mercier, H., Arhan, M., 1989. Mesoscale variability transition from the western to the
eastern Atlantic along 48°N. Journal of Physical Oceanography 19, 1149—1170.
Colin de Verdiere, A., Harvey, J.G., Arhan, M. 1986. Stirring and mixing of thermohaline anomalies.
Journal of Marine Research 44, 93—118.
Cunningham, S.A., Haine, T.W.N., 1995a. Labrador Sea Water in the eastern North Atlantic. Part I:
a synoptic circulation inferred from a minimum in potential vorticity. Journal of Physical Oceanography 25, 649—665.
Cunningham, S.A., Haine, T.W.N., 1995b. Labrador Sea Water in the eastern North Atlantic. Part II:
mixing dynamics and the advective—diffusive balance. Journal of Physical Oceanography 25, 666—678.
Dickson, R., Lazier, J., Meinke, J., Rhines, P., Swift, J., 1996. Progress in Oceanograply 38, 241—295.
Gould, W.J., 1983. The northeast Atlantic Ocean. In: Robinson, A. (Ed.), Eddies in Marine Science.
Springer, Berlin, 609 pp.
Greatbatch, R.L., Goulding, A., 1989. Seasonal variations in a linear barotropic model of the North
Atlantic driven by the Hellerman and Rosenstein wind stress field. Journal of Physical Oceanography
19, 572—595.
Bemer, H.J., Hasse, L., 1987. The Bunker climate Atlas of the North Atlantic Ocean Vol. 2, Air—Sea
Interactions. Springer, Berlin, 218p.
LaCasce, J.H, Speer, K.G., 1998. Langrangian statistics in unforced baratropic flows. Journal of Marine
Research, accepted.
LaCasce, J.H, Speer, 1998. Lagrangian dispersion statistics in the ocean and f/H barotropic control, in
preparation.
Menzel, M., 1998. The Eurofloat experiment. Internal Technical Report No. 98-05, Laboratoire des
Physique des Océans, Ifremer, Brest. 73 pp.
Mercier, H., Colin de Verdiere, A., 1985. Space and time scales of mesoscale motion in the eastern North
Atlantic. Journal of Physical Oceanography 15, 171—183.
Müller, T.J., Siedler, G., 1992. Journal of Marine Research 50, 63—98.
O’Dwyer, J., LaCasce, J., Williams, R., Speer, K., 1998. Large-scale PV conservation observed with floats.
Journal of Physical Oceanography, submitted.
Ollitrault, M., Loaëc, G., Dumortier, C., 1994. MARVOR: a multicycle RAFOS float. Sea Technology 35(2).
Paillet, J., Arhan, M., McCartney, M., 1998. The spreading of Labrador Sea Water in the eastern North
Atlantic. Journal of Geophysical Research 103, 10,223—10,229.
Polzin, K.L., Toole, J.M., Ledwell, J.R., Schmitt, R.W., 1997. Spatial variability of turbulent mixing in the
abyssal ocean. Science 276, 93—96.
Price, J.F., Rossby, H.T., 1982. Observations of a barotropic planetary wave in the western North Atlantic.
Journal of Marine Research 40 (Suppl.), 543—558.
Rhines, P.B., Schopp, R., 1991. The wind-driven circulation: quasi-geostrophic simulations and theory for
nonsymmetric winds. Journal of Physical Oceanography 21, 1438—1469.
Saunders, P., 1982. Circulation in the eastern North Atlantic. Journal of Marine Research 40 (Suppl.),
641—657.
Schmitz, W.J., Hogg, N.G., 1978. Observations of energetic low frequency current fluctuations in the
Charlie Gibbs Fracture Zone. Journal of Marine Research 36, 725—734.
Spall, M.A., 1994. Mechanism for low-frequency variability and salt flux in the Mediterranean salt tongue.
Journal of Geophysical Research 99, 10,121—10,129.
Spall, M.A., Price, J., Richardson, P.H., 1993. Advection and eddy mixing in the Mediterranean salt tongue.
Journal of Marine Research 51, 797—818.
Speer, K.G., Tziperman, E., Feliks, I., 1993. A simple numerical model for topographic and grounding
effects in a bottom layer. Journal of Geophysical Research 98, 8547—8558.
Speer, K.G., 1993. The deep silica tongue in the North Atlantic. Deep-Sea Research 40, 925—936.
Stommel, H., Arons, A.B., 1960. On the abyssal circulation of the world ocean — I. Stationary planetary flow
patterns on a sphere. Deep-Sea Research 6, 140—154.
Sy, A., Rhein, M., Lazier, J.R., Koltermann, K.P., Meinke, J., Putzka, A., Bersch, M., 1997. Surprisingly
rapid spreading of newly formed intermediate waters across the North Atlantic Ocean. Nature 386,
675—679.
K.G. Speer et al. / Deep-Sea Research II 46 (1999) 165—179
179
Talley, McCartney, 1982. Distribution and circulation of Labrador Sea Water. Journal of Physical
Oceanography 12, 1189—1205.
Veronis, G., Stommel, H., 1956. The action of variable wind stresses on a stratified ocean. Journal of Marine
Research 15, 43—75.
Warren, B., 1981. Deep circulation of the world ocean. In: Warren, B.A., Wunsch, C. (Eds.), Evolution of
Physical Oceanography, Scientific Surveys in Honor of Henry Stommel, MIT Press, Cambridge, MA,
pp. 6—41.
Willebrand, J., Philander, S.G.H., Pacanowski, R.C., 1980. The oceanic response to large-scale atmospheric
disturbances. Journal of Physical Oceanography 10, 411—429.
Wilson, C., Speer, K., Charlou, J.-L., Bougault, H., Klinkhammer, G., 1995. Hydrography above the
Mid-Atlantic Ridge in the FAZAR area (33°N—40°N) and within the Lucky Strike segment. Journal of
Geophysical Research 100, 20,555—20,564.
Worthington, L.V., Wright, W.R., 1970. North Atlantic Ocean Atlas of Potential Temperature and Salinity
in the Deep Water Including Temperature, Salinity and Oxygen Profiles from the Erika Dan Cruise of
1962, Woods Hole Oceanographic Institution Atlas Series 2, 24 pp. and 58 plates.