JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 116, C11025, doi:10.1029/2011JC007134, 2011
Vertical structure of mesoscale eddies in the eastern South Pacific
Ocean: A composite analysis from altimetry and Argo
profiling floats
Alexis Chaigneau,1,2,3 Marie Le Texier,4 Gérard Eldin,3 Carmen Grados,2
and Oscar Pizarro5
Received 16 March 2011; revised 7 September 2011; accepted 7 September 2011; published 17 November 2011.
[1] The mean vertical structure of mesoscale eddies in the Peru‐Chile Current System is
investigated by combining the historical records of Argo float profiles and satellite
altimetry data. A composite average of 420 (526) profiles acquired by Argo floats that
surfaced into cyclonic (anticyclonic) mesoscale eddies allowed constructing the mean
three‐dimensional eddy structure of the eastern South Pacific Ocean. Key differences in
their thermohaline vertical structure were revealed. The core of cyclonic eddies (CEs)
is centered at ∼150 m depth within the 25.2–26.0 kg m−3 potential density layer
corresponding to the thermocline. In contrast, the core of the anticyclonic eddies (AEs) is
located below the thermocline at ∼400 m depth impacting the 26.0–26.8 kg m−3 density
layer. This difference was attributed to the mechanisms involved in the eddy formation.
While intrathermocline CEs would be formed by instabilities of the surface equatorward
coastal currents, the subthermocline AEs are likely to be shed by the subsurface
poleward Peru‐Chile Undercurrent. In the eddy core, maximum temperature and salinity
anomalies are of ±1°C and ±0.1, with positive (negative) values for AEs (CEs). This
study also provides new insight into the potential impact of mesoscale eddies for the
cross‐shore transport of heat and salt in the eastern South Pacific. Considering only the
fraction of the water column associated with the fluid trapped within the eddies, each CE
and AE has a typical volume anomaly flux of ∼0.1 Sv and yields to a heat and salt
transport anomaly of ±1–3 × 1011 W and ±3–8 × 103 kg s−1, respectively.
Citation: Chaigneau, A., M. Le Texier, G. Eldin, C. Grados, and O. Pizarro (2011), Vertical structure of mesoscale eddies in the
eastern South Pacific Ocean: A composite analysis from altimetry and Argo profiling floats, J. Geophys. Res., 116, C11025,
doi:10.1029/2011JC007134.
1. Introduction
[2] The Peru‐Chile Current System (PCCS), also known as
the Humboldt Current System, is relatively complex, exhibiting several surface and subsurface currents. Its dynamics is
principally controlled by the atmospheric South Pacific
Anticyclone through Sverdrup dynamics. In the surface layers, the eastward flowing South Pacific Current feeds the
Chile‐Peru Current (CPC) between 30°S and 40°S, forming
the eastern branch of the South Pacific subtropical gyre [Strub
et al., 1998; Chaigneau and Pizarro, 2005a, 2005b]. Near the
South American coast, the persistent equatorward winds
1
Laboratoire d’Océanographie et de Climatologie: Expérimentation et
Analyse Numérique, UMR 7159, IRD, CNRS, UPMC, MNHN, Paris,
France.
2
Instituto del Mar de Perú, Callao, Peru.
3
Laboratoire d’Études en Géophysique et Océanographie Spatiale,
UMR 5566, IRD, CNRS, CNES, UPS, Toulouse, France.
4
Département de Formation en Hydraulique et Mécanique des Fluides,
ENSEEIHT, Toulouse, France.
5
Department of Geophysics, UDEC, Concepcion, Chile.
Copyright 2011 by the American Geophysical Union.
0148‐0227/11/2011JC007134
drive upwelling cells leading to the highest biological productivity of the world ocean in terms of fish [Chavez et al.,
2008]. Over the continental shelf, the upwelling of relatively cold deep water also gives rise to intense thermal fronts
which separate, over short distances, cold coastal water from
warmer and saltier subtropical water of the offshore ocean
[Wyrtki, 1967; Strub et al., 1998]. By geostrophic adjustment,
this relatively strong cross‐shore density gradient in the surface layers reinforces the equatorward Chile‐Coastal Current
and Peru‐Coastal Current (Figure 1a). The PCCS is also
characterized by two major subsurface poleward currents
(Figure 1a): the Peru‐Chile Countercurrent and the Peru‐
Chile Undercurrent (PCU) which have been both traced back
to the equatorial current system [Lukas, 1986; Tsuchiya,
1985; Strub et al., 1998; Montes et al., 2010]. The PCU
which transports relatively warm and salty equatorial subsurface water from the eastern tropical Pacific to at least 48°S
[Silva and Neshyba, 1979] along the continental slope and
shelf is a major source of the coastal upwelling off Peru and
northern Chile [Huyer et al., 1987; Montes et al., 2010;
Albert et al., 2010]. This subsurface current also supplies
low‐oxygenated water that helps maintaining the oxygen
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Figure 1. Spatiotemporal distribution of the 4179 Argo profiles used in this study and typical water mass
properties observed in the PCCS. (a) Position of the 4179 Argo profiles and schematic circulation. Solid
blue lines indicate surface currents, and dashed red lines show subsurface currents. PCC, Peru Coastal
Current; CCC, Chile Coastal Current; PCCC, Peru‐Chile Countercurrent; and PCU, Peru‐Chile Undercurrent. (b) Meridional variation of the number of Argo profiles in 1° latitude bands. (c) Monthly variations of
the number of Argo profiles. (d) Mean temperature‐salinity diagrams observed in the two subregions
delimited by a black dashed line in Figure 1a. Grey curves represent potential density anomalies (s in
kg m−3). SPESTMW, South Pacific Eastern Subtropical Mode Water; ESPIW, Eastern South Pacific Intermediate Water; ESSW, Equatorial Subsurface Water; and AAIW, Antarctic Intermediate Water.
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minimum zone of the eastern South Pacific [Wyrtki, 1963;
Fuenzalida et al., 2009; Stramma et al., 2010].
[3] The oceanic circulation along the South American
coast is also characterized by energetic mesoscale structures,
the oceanic cyclonic and anticyclonic eddies (CEs and AEs,
respectively). Such eddies have been mostly observed from
altimeter data [Chelton et al., 2007; Chaigneau et al., 2008,
2009], but they also have a signature in color satellite
images [Correa‐Ramirez et al., 2007] or in surface‐drifter
trajectories [Chaigneau and Pizarro, 2005c]. In the PCCS,
CEs and AEs, which have a typical diameter of 150–
300 km, are principally formed near the South American
coast where they locally impact the heat and salt budgets
through lateral turbulent fluxes [Chaigneau and Pizarro,
2005b; Colbo and Weller, 2007]. Then, CEs and AEs,
which acquire a water mass structure typical of their formation region, propagate seaward with translation velocities of
few cm s−1 owing to a combination of mean flow advection
and self‐propagation. In this new environment, eddies
appear as anomalous water masses with surface or subsurface temperature and salinity anomalies [Johnson and
McTaggart, 2010] which are progressively redistributed to
surrounding water during eddy decaying phase [e.g., Swart
et al., 2008]. The westward propagation of CEs can also
extend the area of high biological productivity offshore by
both eddy chlorophyll advection and eddy nutrient pumping
[Correa‐Ramirez et al., 2007].
[4] Although the main horizontal structure and kinematic
properties of the PCCS eddies were investigated during the
last decade, very little is known about their vertical structure
and their impact on the heat and salt transports. Chaigneau
and Pizarro [2005c] have briefly examined the vertical
structure of a particular CE sampled by the World Ocean
Circulation Experiment (WOCE) P19 hydrographic section
along 88°W, showing that the subsurface CE core have typical temperature and salinity anomalies of around −1°C and
−0.1, respectively. More recently, Johnson and McTaggart
[2010] used Argo float profile data to characterize AEs
of the PCCS, showing that their core is located in the
subthermocline and contains anomalous signature of the
Equatorial Subsurface Water originating from the PCU.
Although this analysis provides a first insight into the
vertical characteristics of AEs in the PCCS, some open
questions remain. For instance, do CEs exhibit the same
vertical structure as AEs? What is the respective role of such
eddies on the volume, heat and salt transports?
[5] The main objective of this work is thus to answer these
questions and extend the work of Johnson and McTaggart
[2010]. To achieve this goal, we propose a blended analysis of satellite altimetry data and Argo float profiles allowing
the three‐dimensional reconstruction of composite eddies.
The obtained average eddy fields are used to characterize the
mean vertical structure of both eddy types and to estimate
their contribution to the volume, heat and salt transports. The
paper is organized as follows. In section 2 we describe the
study region and the two data sets (altimetry and Argo data)
used in this work. Eddy identification algorithm from satellite data is also briefly presented in this section as well as the
methodology used to classify the Argo profiles. The methodology used for the construction of the three‐dimensional
composite eddies is presented in section 3. The mean thermohaline properties of the composite CE and AE are
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described in section 4, as well as the eddy function in the
offshore transport of volume, heat and salt. Section 4 also
deals with the meridional variation of the eddy vertical
structure between the northern and southern areas of the
study region. The observed differences between cyclonic and
anticyclonic eddies allow a discussion in section 5 of the
most probable origin of both eddy types. Finally in section 6,
we summarize the results.
2. Data and Methods
[6] The study region extends from 10°S to 30°S and from
the South American coast to 100°W (Figure 1a). The
northern boundary is located south of the equatorial front
which separates the relatively fresh equatorial water from
the saltier subtropical water [Fiedler and Talley, 2006].
Similarly, the southern margin is located to the north of the
subtropical front, separating the subtropical water from the
relatively cold and fresh subantarctic water [Tomczak and
Godfrey, 1994; Stramma et al., 1995; Chaigneau and
Pizarro, 2005a].
2.1. Altimeter‐Derived Sea Level Anomaly Data
and Eddy Identification
[7] The presence and position of mesoscale eddies in the
study domain are determined by analyzing sea level
anomaly (SLA) maps from the multisatellite AVISO product
(http://www.aviso.oceanobs.com) between December 2003
and October 2009. This gridded multimission altimeter
product, produced by Ssalto/Duacs and distributed by CLS–
Space Oceanography Division (Toulouse, France), provides
the best available spatiotemporal resolution for observing
mesoscale features [Le Traon and Dibarboure, 1999; Pascual
et al., 2006]. SLA data, computed relatively to a 7 year mean
(1993–1999), are weekly mapped onto a 1/3° Mercator grid
and then bilinearly interpolated by AVISO onto a 0.25° ×
0.25° longitude/latitude grid (see Appendix A.2 of Chelton
et al. [2011] for a detailed description of the procedure
used by AVISO to produce SLA maps). This weekly SLA
product is further linearly interpolated in time to obtain SLA
maps on a 1 day basis.
[8] In each daily SLA map, mesoscale eddies were identified using the method recently developed by Chaigneau
et al. [2009]. The CE (AE) detection algorithm first
involves searching for eddy centers associated with local
SLA minima (maxima) in a moving window of 1° × 1°. Then,
for each possible cyclonic (anticyclonic) center, the algorithm
looks for closed SLA contours with an increment (decrement)
of 10−3 m. The outermost closed SLA contour, embedding
only the considered center, corresponds to the eddy edge.
Given the accuracy of satellite measurements and AVISO
product [Le Traon and Ogor, 1998; Ducet et al., 2000;
Chelton and Schlax, 2003], we only consider here eddies
having SLA amplitudes higher than 2 cm.
[9] A similar algorithm, based on closed SLA contours, has
also been used recently to describe mesoscale eddy characteristics at global scales [Chelton et al., 2011]. This eddy
identification method has been shown to be more accurate
than the commonly used Okubo Weiss‐based method [e.g.,
Isern Fontanet et al., 2003]. In particular, it allows to considerably reduce the number of “false” detected eddies
[Chaigneau et al., 2008]. Souza et al. [2011] have also
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shown, through the comparison of three automatic detection
algorithms in the South Atlantic, that the geometric criterion
used in this study exhibits a better performance, mainly in
terms of the eddy detection number, their lifetimes and
propagation velocities.
2.2. Argo Data Set
[10] The vertical structure of mesoscale eddies was
investigated using the autonomous CTD profiling floats of
the Argo program (http://www.coriolis.eu.org) available in
the study region. This data set spans the same temporal
period as the altimetry product (December 2003 to October
2009). The original data set consists of 93 distinct floats
which acquired 6050 CTD profiles during the studied period.
Argo data are automatically preprocessed and quality controlled by the Argo data centers [Wong et al., 2003; Böhme
and Send, 2005; Owens and Wong, 2009]. Only pressure
(P), temperature (T), and salinity (S) data flagged as good
(Argo quality flag 1) were retained in the analyses. However,
some of these preprocessed profiles were still suspicious.
Thus, for our analysis we only considered profiles for which
(1) the shallowest data are located between the surface and
10 m depth and the deepest acquisition is below 1000 m;
(2) the depth difference between two consecutive data does
not exceed a given limit (Dzlim) which depends on the
considered depth (Dzlim = 25 m for the 0–100 m layer;
Dzlim = 50 m for the 100–300 m layer; and Dzlim = 100 m
for the 300–1000 m layer); and (3) the number of data levels
in the 0–1000 m layer is higher than 30.
[11] Once selected, every profile was visually checked,
and those which presented a suspicious T/S diagram were
systematically discarded. The final database is composed of
4179 profiles, corresponding to 69% of the initial data set
(Figure 1a). The PCCS was rather homogeneously sampled
by Argo floats except around 25°S where fewer profiles
were acquired (Figure 1b). The temporal distribution shows
that the number of CTD profiles linearly increased from
December 2003 (date of the first Argo float deployment in
the study region) to June 2008 (Figure 1c) and dramatically
decreased afterward.
[12] The mean T/S diagrams for the northern (10°S–20°S)
and southern (20°S–30°S) regions of the study domain
highlight the main water masses present in the PCCS
(Figure 1d). South Pacific Eastern Subtropical Mode Water
fills the major part of the surface ocean, with T > 20°C and
S > 35. In subsurface, the relatively fresh and cold Eastern
South Pacific Intermediate Water (T = 12–13°C, S < 34.8,
s ≈ 26 kg m−3) is located at the base of the thermocline and
surmounts the colder and saltier Equatorial Subsurface
Water (T ≈ 8–11°C, s ≈ 26.5 kg m−3). Finally, below the
Equatorial Subsurface Water is found the Antarctic Intermediate Water associated with a local salinity minimum
(T < 7°C, S < 34.6, s > 27 kg m−3). Note that both the
Eastern South Pacific Intermediate Water and the Antarctic
Intermediate Water become saltier and slightly warmer along
their northward/northwestward pathway in the PCCS,
whereas the Equatorial Subsurface Water of equatorial origin
gets fresher and colder toward the south (Figure 1d). Readers
interested in more details on the water mass distribution in
the eastern South Pacific should refer for instance to Wyrtki
[1967], Strub et al. [1998], Blanco et al. [2001], Wong and
Johnson [2003], and Schneider et al. [2003].
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[13] The 4179 retained T/S profiles were then linearly
interpolated onto 101 regularly spaced vertical levels every
10 m from the surface to 1000 m, assuming that the shallowest values of T and S were representative of surface
values. Derived quantities were estimated from these vertically gridded values of T, S, such as potential temperature
() and density (s). At each vertical level, the dynamic
height (DH) relative to the 1000 m reference depth was also
computed. We considered that a 1000 m level of no motion
should not strongly impact our results since the average flow
field at 1000 m depth, estimated from Argo drift data has a
magnitude weaker than 1 cm s−1 in the PCCS [Davis, 2005;
Katsumata and Yoshinari, 2010].
[14] Finally, to better describe the mesoscale perturbations
captured by the floats, anomalies of every property (′, S′,
s′ , DH′) were computed by removing climatological profiles. Here, the climatological T/S profiles were obtained by
interpolating the CSIRO Atlas of Regional Seas (CARS)
[Dunn and Ridgway, 2002; Ridgway et al., 2002] to floats’
positions and times. Computed anomalies are expected to be
relatively weak for CTD profiles acquired outside eddies
and relatively strong inside eddies.
2.3. Classification of the Argo Profiles
[15] The 4179 CTD profiles were classified among three
distinct categories depending whether the floats surfaced
outside an eddy (OE) or inside a CE or AE identified from
daily SLA maps. Figure 2a shows the result of the classification for a given day in November 2005 with two floats
surfacing in CEs, one in an AE and two OEs. Over the entire
study period (December 2003 to October 2009), we obtained
a total of 420 profiles inside CEs, 526 inside AEs and
3233 OEs (Figures 2c–2e). Thus, AEs were slightly more
sampled than CEs and ∼23% of the profiles surfaced into
an eddy. This proportion is in agreement with the results of
Chaigneau et al. [2008], who estimated that mesoscale
structures cover ∼25% of the PCCS area.
[16] It is important to note that, apart from potential errors
in the automated eddy‐edge identification, three main factors
could cause errors in the classification of the Argo floats.
First, the accuracy of the CTD profiles locations depends on
the precision of the Argos positioning system which varies
from 150 m for a Class 3 Argos location to 1000 m for a
Class 1 Argos location. Second, the exact surfacing position
assigned to a CTD profile cannot be determined accurately
since an Argos positioning satellite does not necessarily fly
over a float immediately after it surfaces. A typical lag of
1 hour, during which Argo floats are advected by surface
currents, can be expected between the exact surfacing
location and the first reported position [Lebedev et al.,
2007]. However, considering maximum surface currents of
∼10 cm s−1 in the PCCS [Chaigneau and Pizarro, 2005b],
the error between reported CTD position and surfacing
position is relatively small (<500 m). Third, eddies were
identified from daily interpolated SLA maps and not at the
exact surfacing times reported by the Argo floats. On average, the 4179 surfacing times were recorded 1.3 h ahead the
associated SLA maps. Then, since long‐lived eddies propagate westward with typical velocities of 2–7 cm s−1 in the
study region [Chaigneau and Pizarro, 2005c; Chaigneau
et al., 2009], an additional small error of 100–300 m is
considered. We estimate that the total error introduced by
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Figure 2. An illustrative example of the eddy detection algorithm and classification of the Argo profiles.
(a) Color shading corresponds to sea level anomalies (in cm) for the map of 15 November 2005, whereas the
edges of the automatically identified cyclonic and anticyclonic eddies are indicated by blue and red contours, respectively; Argo floats that surfaced on 15 November 2005 are shown as black dots. (b) Illustration
of the floats’ positions (M1 and M2) relative to the corresponding eddy centers (C1 and C2); this eddy‐
centered referential (Dx, Dy) is used to construct the composite eddies through objective interpolation
(see Figure 3 and text for details). (c) Position of the 420 Argo profiles that surfaced inside cyclonic
eddies between December 2003 and October 2009. (d) Same as Figure 2c but for the 526 profiles surfacing
inside anticyclonic eddies. (e) Same as Figure 2c but for the 3233 profiles surfacing outside eddies.
synchronization biases and positioning errors is of ∼1 km
on average, but can probably reach a distance as large as
∼5 km. This error can result in an erroneous classification of
a few Argo floats, in particular when they are close to the
identified eddy edges. However, only 0.1% (3.6%, respectively) of the CTD profile positions is closer than 1 km
(5 km) from a detected eddy edge. It is also worth to note that
the uncertainty of the eddy center and edge location determined from the SLA maps are probably much larger than the
synchronization biases and positioning errors of the Argo
floats.
3. Construction of Cyclonic and Anticyclonic
Composite Eddies
3.1. Objective Mapping of Eddy Properties
[17] CTD profiles surfacing into CEs or AEs can be
localized relative to their corresponding eddy centers identified as local maxima/minima in SLA (Figure 2b). Assuming first that all the CEs or AEs of the PCCS exhibit similar
3‐D structures, the spatial distribution of a given property (or
anomaly) inside CEs and AEs is thus investigated by a
composite analysis using a coordinate system (Dx, Dy) in
which each CTD profile is located respectively to the corresponding eddy center (Dx = Dy = 0). For instance, solid
dots in Figures 3a–3d show the temperature and salinity
anomaly distributions as a function of the distance (Dx, Dy)
to the eddy center. To illustrate the method, we chose a level
of 150 m for CEs (Figures 3a, 3c, and 3e) and 400 m for AEs
(Figures 3b, 3d, and 3f) which correspond to the vertical
location of the eddy cores (see section 4).
[18] The profiles inside the CEs and AEs are rather
homogeneously distributed around the composite eddy center. For example, at 150 m inside the CEs, the mean temperature and salinity anomalies obtained from the selected
420 Argo profiles are of −0.40°C and −0.05, respectively
(Table 1). At 400 m inside the AEs, the mean values obtained
are of 0.48°C and 0.04, respectively (Table 1). Although the
variability around these mean values is large, around 70%
(80%, respectively) of the 420 (526) ′ and S′ values are
negative (positive) for CEs (AEs). The mean ′ and S′ values
and the percentage of negative/positive values are strongly
different from the ones computed from the 3233 profiles
located OEs (Table 1). The nonparametric Mann‐Whitney
U‐test used to compare the different anomaly distributions
[Scherrer, 2007], confirmed that anomalies computed inside
eddies were significantly different (p < 0.05) from the ones
obtained OEs. Unlike the parametric t test, this nonparametric test makes no assumptions about the distribution of
the data (e.g., normality and equality of variance).
[19] At each depth level, anomalies that are more than
3 times the interquartile range from either the first or third
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Figure 3. Objective interpolation, onto a regular grid (10 × 10 km2) of (a, b) potential temperature
anomalies and (c, d) salinity anomalies. Anomalies are shown at 150 m depth for cyclonic eddies
(Figures 3a and 3c) and 400 m depth for anticyclonic (Figures 3b and 3d) eddies. Solid dots in
Figures 3a–3d represent the anomalies estimated from Argo profiles in the eddy‐centered referential,
whereas color shadings correspond to the results of the objective interpolation. (e, f) Objectively interpolated dynamic height anomaly at 150 and 400 m depths relative to 1000 m (black contours) and horizontal
geostrophic speed (color shading, in cm s−1), respectively. White dots in Figures 3e and 3f show composite
eddy edges identified from the automatic algorithm, whereas black lines show eddy edges after fitting to an
ellipse (see text for details). Anomalies were computed relatively to the CSIRO Atlas of Regional Seas
(CARS) climatology interpolated to the positions and times of the Argo floats.
quartiles were considered as outliers and were discarded. The
remaining properties were then objectively mapped, assuming an isotropic Gaussian covariance decorrelation scale of
100 km. This scale allows filtering out the undesired small‐
scale variability and approximately corresponds to the mean
eddy radius previously observed in the study region
[Chaigneau et al., 2008, 2009]. Indeed from altimetry data,
the mean radius of the 971 eddies sampled by the profiling
floats is of 122 ± 30 km. Using the objective interpolation,
properties were mapped onto a regular 10 × 10 km2 grid.
This grid spacing was chosen since in the new referential,
80% of the minimum distances between two profiles are less
than 10 km (Figures 3a–3d). Color shading in Figures 3a–3d
shows the objectively mapped ′ and S′ at 150 m and 400 m
depth. At these levels, maximum ′ of 0.5°C–1°C are
observed near the composite center at Dx = Dy = 0. Although
negative (positive) anomalies persist over the interpolation
domain for CEs (AEs), the anomalies strongly weaken at
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Table 1. Seawater Property Anomalies ′ and S′ Obtained From the 420 (526) Argo Floats That Surfaced Inside CEs (AEs)a
150 m Depth
N
′ (°C)
′ /max
′ (°C)
min
S′
′ /Smax
′
Smin
400 m Depth
Cyclonic Eddy
Outside Eddies
Anticyclonic Eddy
Outside Eddies
420
−0.40 ± 0.67
−2.96/1.88 (75%)
−0.05 ± 0.13
−0.61/0.35 (63%)
3233
0.10 ± 0.73
−2.75/2.85 (46%)
0.01 ± 0.14
−0.51/0.54 (48%)
526
0.48 ± 0.55
−0.69/2.58 (84%)
0.03 ± 0.05
−0.09/0.22 (76%)
3233
0.04 ± 0.33
−1.18/1.36 (53%)
0.00 ± 0.03
−0.12/0.13 (47%)
a
These values are compared to those obtained from the 3233 floats located outside eddies. A depth of 150 m (400 m) was chosen for CEs (AEs) since it
corresponds to the depth where density anomalies s′ are maximum. Numbers for ′ and S′ indicate the average ±1 standard deviation. Percentages
indicated correspond to the percentage of negative (positive) values observed at 150 m (400 m) depth in CEs (AEs) and OEs.
50–100 km from the eddy center. Thus at these levels, the
eddy core extends less than 100 km from the eddy center.
At 150 m depth in the core of the CE, the maximum
salinity anomalies are centered slightly more northward
than the associated temperature anomalies (Figure 3c). This
difference is induced by the objective mapping applied to
the much more variable salinity field (Figure 3c).
3.2. Horizontal Extent of the Composite Eddy Cores
[20] Figures 3e and 3f show the composite DH′ (black
contours) at 150 m and 400 m depth associated with CEs and
AEs and the corresponding geostrophic velocities computed
from the DH′ slopes (quivers and color shading). At each
vertical level, the center of the composite eddies correspond
to the local minimum/maximum of DH′ closest to the grid
center (Figures 3e and 3f). To determine the horizontal extent
of the eddy core, we averaged the geostrophic swirl velocities
along closed DH′ contours embedding the eddy center. The
closed DH′ contour associated with the strongest average
swirl velocity corresponds to the composite eddy‐core edge
(white dots in Figures 3e and 3f). To avoid spurious shapes
and vertical discontinuity, an ellipse is then least squares
fitted on this contour (black heavy ellipse in Figures 3e
and 3f). By definition, the composite eddy‐core edge coincides with a local maximum in geostrophic speed. This
maximum value can be highly variable from eddy to eddy
since it is proportional to eddy amplitudes for similar radii.
At 150 m depth (400 m, respectively), the maximum swirl
velocity value is of ∼7.8 cm s−1 (6.9 cm s−1) for the composite CE (AE). At these depths and considering only the
profiles located inside the ellipse, the mean temperature and
salinity anomalies of the eddy cores are on the order of
±0.8°C and ±0.05, respectively.
3.3. Vertical Extent of the Trapped Fluid
[21] One of the main objectives of this study is to estimate
the volume, heat and salt transports associated with the
PCCS eddies (see section 4.3). However, the water mass
anomalies in an eddy can only be maintained if the water
mass in the eddy is trapped for a considerable part, preventing surrounding water to enter the eddy when the eddy
is translated [Flierl, 1981; van Aken et al., 2003]. Thus, in
order to not overestimate the lateral transports we must only
consider the part of the water column which is effectively
trapped and transported by the eddies. On the basis of the
suggestions made by Flierl [1981], the amount of water
trapped inside a ring depends on the ratio of its drift velocity
to its tangential velocity. This ratio, which arises from the
comparison of nonlinear advection and acceleration, pro-
vides a measure of nonlinearity stating that the eddy
dynamics is nonlinear when this ratio exceeds 1 and maintains a coherent structure as it propagates [e.g., Flierl, 1981;
Chelton et al., 2007, 2011]. Thus, the composite eddies
were tracked with depth as long as their rotational speed
exceeded their translation speed. The rotational speeds of
our composite eddies were calculated at each depth by
averaging the geostrophic velocity along the eddy‐core
edge, and we defined their translation speed as being 4.3 cm
s−1. This value corresponds to the mean propagation speed
of long‐lived eddies in the study region [Chaigneau et al.,
2009]. Figure 4 shows the vertical profiles of the mean
swirl velocity at the eddy‐core edge for both the CE and AE.
The vertical extent of the trapped fluid in the composite CE
is of 240 m and of 530 m in the AE. Again, these values are
probably variable eddy to eddy and might be proportional to
the amplitude of the individual eddies. Hereafter, the term
“eddy” will refer to the eddy core extending from the surface to these “trapping depths.”
[22] Table 2 shows that on average along the vertical, the
composite AE is slightly smaller and of a more circular
shape than the composite CE. The mean equivalent radius
is of ∼60 km for both composite eddies corresponding to
typical areas of ∼10000 km2. These spatial scales correspond to the radii of the eddy cores which extend from the
eddy center to the closed DH′ contour associated with a
maximum rotational speed. The size of the eddy cores is
about half the total eddy size of 122 km determined previously from altimetry (see section 3.1). Table 2 also suggests
that both composites AE and CE are not perfectly circular
but slightly elongated along a southwestward/northeastward
direction. The mean ratios of the major to the minor axes of
the ellipses are of 1.3 for CE and 1.2 for AE, in agreement
with the values of 1.5–1.7 obtained at the surface from
altimetry [Chaigneau et al., 2008].
4. Eddy Vertical Structure and Associated
Transports
4.1. Mean Vertical Thermohaline Anomalies
of Composite Eddies
[23] In order to illustrate the differences between CE and
AE in terms of thermohaline structure, Figure 5 shows
vertical sections of the composite eddies along the zonal
direction at Dy = 0. CE (AE, respectively) shows a maximum
temperature anomaly of about −1°C (+1°C) centered at 150 m
(400 m). Although temperature anomalies greater than 0.5°C
extend between 50 m and 350 m for CE (Figure 5a) and
between 200 m and 600 m for AE (Figure 5b), we still observe
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Figure 4. Vertical profiles of the swirl velocity averaged over the composite cyclonic (blue) and anticyclonic (red) eddy edges. The lower scale corresponds to the swirl velocity (in cm s−1), whereas the
upper scale corresponds to the nonlinearity parameter (ratio of the swirl velocity to the drift velocity)
using a mean drift velocity of 4.3 cm s−1. The indicated depths correspond to the vertical extent of the
trapped fluid within the composite eddies (240 m for CE and 530 m for AE).
a weak temperature anomaly of ±0.15°C at 1000 m depth.
Similarly, maximum salinity anomalies of ±0.1 are approximately centered at the same vertical levels than their temperature counterparts (Figures 5c and 5d). For the composite
AE, we note, however, that (1) maximum salinity anomalies
are slightly shifted upward compared to temperature
anomalies and (2) the eddy core is surmounted by relatively
fresh and cold water. This latter feature, also noted in model
simulations off Peru [Colas et al., 2011], is probably due to
the doming of the isotherms above the AE cores. The negative
temperature anomalies in the upper layers also suggest that
sea surface temperature alone would not be a good variable
for detection of AEs.
[24] Along the zonal section at Dy = 0, the combination of
′ and S′ leads to meridional geostrophic velocity anomalies
on the order of ±10 cm s−1, with maximum anomalies near
eddy‐core edges (Figures 5e and 5f). These geostrophic
velocities were computed considering the Coriolis parameter
at 20°S which corresponds to the mean latitude of the study
region. Again, there is a notable difference in terms of eddy
kinematics, since maximum velocity anomalies are observed
in the upper 0–200 m layer for CE and between 150 m and
400 m depth for AE (see also Figure 4). At the surface,
geostrophic velocities vary from a few cm s−1 near the
center to 7–10 cm s−1 at the eddy edges (Figures 4 and 5).
These later values are consistent with the eddy swirl speed
of 10–12 cm s−1 estimated from near‐surface drifter trajectories and altimetry measurements [Chaigneau and Pizarro,
2005c].
[25] Figures 6a–6c show the mean vertical temperature,
salinity and density anomalies obtained inside the ellipses
determining the composite eddy edges. These vertical
profiles are also compared with those obtained by averaging
all CTD profiles located either inside CEs, or inside AEs, or
OEs. In the eddy core of the composite eddies, we observed
mean maximum anomalies larger than ±0.7°C in temperature and ±0.06 in salinity. These /S anomalies lead to
maximum density anomalies on the order of +0.10 kg m−3
(−0.08 kg m−3, respectively) for CE (AE). Again, we note
the three main characteristics previously observed along the
zonal vertical section: (1) a core with maximum anomalies
centered at around 150 m for CE and 400 m depth for AE,
(2) a maximum salinity anomaly centered slightly above the
maximum temperature anomaly for AE, and (3) a layer of
relatively fresh and cold water overlaying the AE core. Thus,
Figures 5 and 6 show that on average, the thermohaline
structure of composite CEs and AEs and associated kinematics strongly differs. The core of the composite CE is centered at 150 m depth, a level corresponding to the seasonal
Table 2. Mean Geometrical Properties of the Composite Cyclonic
and Anticyclonic Eddiesa
Vertical extent (m)
Semimajor axis A (km)
Semiminor axis B (km)
Angle 8 (deg)
A/B
Radius (km)
Area (× 103 km2)
Cyclonic Eddy
Anticyclonic Eddy
240
72.3 ± 3.4
53.7 ± 2.3
14.1 ± 5.6
1.3 ± 0.0
62.3 ± 2.8
12.2 ± 1.1
530
63.4 ± 1.6
52.4 ± 1.2
44.7 ± 8.4
1.2 ± 0.0
57.6 ± 1.2
10.4 ± 0.4
a
Numbers indicate the vertical average ±1 standard deviation. A and B
represent the semimajor and semiminor axes of the fitted ellipses,
whereas 8 corresponds to the ellipse orientation relative to the eastward
direction.
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Figure 5. Vertical section at Dy = 0 across the composite cyclonic and anticyclonic eddies. (a, b) Potential temperature anomaly (°C), (c, d) salinity anomaly (× 10−2), and (e, f) meridional (cross section) geostrophic speed (cm s−1) relative to 1000 m, indicating the clockwise (anticlockwise, respectively) rotation
for the composite cyclonic (anticyclonic) eddy. Eddy edges are denoted by black lines, whereas horizontal
dashed lines indicate the trapping depths.
thermocline/halocline/pycnocline characterized by a temperature of ∼14°C, a salinity of ∼34.7 and a density of
∼26 kg m−3 (Table 3). In contrast, the core of the composite
AE is centered below the seasonal thermocline at 400 m
depth, a level associated with Equatorial Subsurface Water
with a temperature of 9.5°C, a salinity of ∼34.6 and a
density of ∼26.7 kg m−3. The origin of this important difference between CE and AE morphology will be discussed
in section 5.
[26] Composite anomalies are consistent with those
obtained from all CTD profiles (thin lines in Figures 6a–6c)
but are stronger since they only correspond to the eddy core
located inside the ellipse. Outside the composite eddy core
and far from the eddy center, anomalies are weaker and
hence the average vertical profiles determined directly from
CTD measurements show weaker anomalies. Although
temperature and density anomalies are significantly different
from anomalies obtained outside eddies down to 1000 m
depth (Figures 6a and 6c; see also Figure 5), the fluid which
is more likely trapped and transported by the composite
eddies is limited to the upper 240 m for CE and 530 m for
AE. Table 4 shows the mean anomalies estimated from the
Argo profiles located within the trapped fluid. It confirms
that anomalies are relatively similar but of opposite signs
between CE and AE and are much larger than outside
eddies. Finally, as suggested by Figure 6 and Table 4, CTD
profiles located OEs show on average small positive biases
of 0.05°C in temperature and 0.5 × 10−2 in salinity compared to the CARS climatology. The positive salinity bias is
relatively strong in the upper 250 m and almost 0 in the
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Figure 6. Mean vertical profiles of (a) temperature anomaly, (b) salinity anomaly, and (c) density
anomaly inside eddies. Thick lines correspond to the mean profiles obtained inside the composite CE (blue
line) and AE (red line). Thin lines correspond to the mean profiles obtained using all the Argo profiles
located inside the 420 CEs (blue line) and 526 AEs (red line). Black thick lines correspond to the mean
vertical profiles obtained from the 3233 Argo profilers located outside the eddies. Anomalies were computed relatively to the CARS climatology interpolated to the positions and times of the Argo floats.
deepest layers (black curve in Figure 6b). In contrast, the
positive temperature bias is observed from the surface to
1000 m depth (black curve in Figure 6a).
4.2. Meridional Variation of the Eddy Vertical
Structure
[27] In order to estimate the robustness of the previous
results and to study the potential meridional variation in the
eddy vertical structure, the same composite analysis was
repeated for the northern (10°S–20°S) and southern (20°S–
30°S) regions of the study domain (see Figure 1a). The
northern region includes 210–220 CTD profiles in both CEs
and AEs whereas the southern region contains a similar
number of profiles in CEs but 302 in AEs. In both regions,
the radius of composite eddies are similar and of ∼58 km
for AEs and ∼61 km for CEs.
[28] Figure 7 shows the eddy vertical structure along the
zonal direction at Dy = 0 for both regions. Vertical structures of eddies in the two regions are qualitatively similar to
the mean structure obtained for the whole PCCS. However,
Figure 7 shows that temperature and salinity anomalies
associated with CEs are larger in the northern region. In
contrast, the composite AE of the northern region shows
smaller ′ and S′ than in the southern region. The swirl
velocity of the composite eddies are maximum in the surface
for CEs (Figures 7i and 7k) and in subsurface for AEs
(Figures 7j and 7l), but eddies rotate more rapidly in the
northern region.
[29] Another clear difference between the northern and
southern regions is the vertical position of the eddy cores.
For both AEs and CEs the maximum ′ and S′ are shallower
in the northern region. For instance, maximum temperature
anomalies are found at 100 m depth in the composite CE of
the northern region and at 150 m depth in the southern
region (Figures 7a and 7c). Similarly, maximum ′ are found
at 350 m depth in the northern composite AE and at ∼500 m
depth in the southern AE (Figures 7b and 7d). Again, the
salinity anomaly cores in the AEs are found slightly above
the maximum ′ levels. Independently of the considered
subregion and as also noted in Figure 5, the relatively warm
and salty AE cores are overlaid by relatively fresh and cold
water. Finally, the vertical extent of the eddy core transporting trapped fluid can be estimated considering a mean
drift velocity of 5.6 cm s−1 between 10°S–20°S and of
3.2 cm s−1 between 20°S–30°S [Chaigneau et al., 2009].
For CEs, this vertical extent is of 200 m in the North and
280 m in the South. For AEs, the trapping depth also
increases southward, varying from 450 m in the northern
region to 570 m in the southern region.
4.3. Volume, Heat, and Salt Transports
[30] In sections 4.1 and 4.2, the combination of altimetry
data and Argo profiles has been shown to be ideally suited to
produce a good approximation of the three‐dimensional
structure of both cyclonic and anticyclonic eddies. We thus
now use this three‐dimensional structure to estimate the relative eddy contribution to fluxes of volume, heat, and salt in
the PCCS. In general, annual mean “eddy” transports can be
estimated by dividing the volume, heat or salt anomaly of one
eddy by 1 year and taking into account the number of eddies
per year [e.g., Gordon and Haxby, 1990; van Ballegooyen
et al., 1994; Doglioli et al., 2007]. The underlying assumption is that mesoscale eddies are sufficiently nonlinear so that
Table 3. Mean Seawater Properties (, S, s) and Their Anomalies
(′, S′, s′) Observed in the Core of the Composite Eddies at the
Depths Where Density Anomalies s′ Are Maximuma
Depth (m)
(°C)
S
s (kg m−3)
′ (°C)
S′
s′ (kg m−3)
Cyclonic Eddy
Anticyclonic Eddy
150
13.63 ± 0.21
34.68 ± 0.06
25.99 ± 0.05
−0.72 ± 0.14
0.05 ± 0.01
0.11 ± 0.03
400
9.53 ± 0.12
34.61 ± 0.02
26.73 ± 0.01
0.77 ± 0.10
0.06 ± 0.01
−0.08 ± 0.01
a
Here , S, and s are mean seawater properties and ′, S′, and s′ are
their anomalies. Density anomalies s′ are maximum at 150 m for CE
and at 400 m for AE. These values were obtained directly from the
objectively interpolated gridded fields. Numbers indicate the average
±1 standard deviation.
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Table 4. Mean Temperature, Salinity, and Density Anomalies of the Eddy Cores From the Surface Down to the Trapping Depths of
240 m for CEs and 530 m for AEsa
Mean ′ (°C)
Mean S′ (× 10−2)
Mean s′ (× 10−2 kg m3)
Cyclonic Eddies
(0–240 m)
Outside Eddies
(0–240 m)
Anticyclonic Eddies
(0–530 m)
Outside Eddies
(0–530 m)
−0.47 ± 0.25
−3.7 ± 2.4
6.8 ± 3.9
0.07 ± 0.04
0.6 ± 0.5
−1.5 ± 0.6
0.41 ± 0.40
4.5 ± 3.8
−3.3 ± 4.9
0.06 ± 0.03
0.3 ± 0.4
−1.2 ± 0.5
a
Values were obtained considering only the Argo profiles located in the core of the composite eddies (93 inside CEs and 162 inside AEs) and can be
compared to the values obtained outside eddies in the same vertical ranges. Numbers indicate the vertical average ±1 standard deviation.
the anomalies are trapped inside their core, down to the
“trapping depth” defined using the Flierl [1981] criterion.
[31] On average over the PCCS, the volume of trapped
fluid transported by an eddy is estimated to be 3.4 × 1012 m3
for a CE and 5.1 × 1012 m3 for an AE (Table 5). Spread over
a 1 year period, the westward volume flux associated with
the trapped fluid of a single eddy is of 0.1–0.2 Sv (Table 5).
Table 5 also indicates that these volume transports are ∼40%
weaker in the northern region than in the south, owing to a
shallower trapping depth and thus a reduced volume. To
estimate how much warm and salty (cold and fresh,
respectively) water is being transported by AEs (CEs), we
calculated the available heat and salt content anomalies
(AHA and ASA) per meter on the vertical:
Z
AHA ¼
Cp ′dA;
ð1Þ
Z
ASA ¼ 0:001
S′dA;
ð2Þ
where r is the density (in kg m−3), Cp is the specific heat
capacity (4000 J kg−1 K−1), and ′, S′ are integrated over the
area (A) of the composite eddy delimited by the ellipse. The
Figure 7. Same as Figure 5 but for the northern and southern regions delimited by the black dashed line
in Figure 1a. The trapping depths related to the vertical extent of the trapped fluid were estimated using a
drift velocity of 5.6 and 3.2 cm s−1 in the northern and southern regions, respectively.
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Figure 8. Mean available (a) heat anomaly and (b) salt anomaly, inside composite CE (blue lines) and
AE (red lines). Horizontal dashed lines indicate the trapping depths.
factor 0.001 converts salinity to salinity fraction (kg of salt
per kg of seawater). Figures 8a and 8b show the vertical
profiles of AHA and ASA within composite CE and AE. In
the core of the composite eddies, maximum AHA are larger
than ±0.3 × 1018 J m−1 whereas ASA are larger than ±0.8 ×
1010 kg m−1. We also note a secondary local maximum of
AHA and ASA at ∼300 m depth in CE. This local maximum, which is related to a larger eddy area in the deeper
layers (see Figure 5), is however probably not trapped and
thus not transported by the eddies.
[32] Integrated from the surface to the trapping depth
(240 m for CE and 530 m for AE), the total available heat
and salt content anomalies transported by a single CE (AE,
respectively) is of −5.5 × 1018 J (+8.7 × 1018 J) and −9.8 ×
1010 kg (+23.8 × 1010 kg) (Table 5). Spread over a 1 year
period, the anomalies of heat and salt transports associated
with the trapped fluid of a single CE (AE, respectively) are
of −1.7 × 1011 W (+2.8 × 1011 W) and −3.1 × 103 kg s−1
(+7.5 × 103 kg s−1), respectively (Table 5). The heat and
salt transports associated with CEs (AEs) are stronger
(weaker) in the northern region than in the southern region.
[33] Applying the eddy tracking algorithm developed by
Chaigneau et al. [2009] on the SLA data set, we determined
that around 310 eddies are formed each year near the coast
between 10°S and 30°S (not shown). Generally, these
eddies are quickly dissipated and thus only influence near‐
coastal regions. However, among these 310 eddies generated each year around 30 CEs and 25 AEs can be tracked for
more than 3 months and propagate at a mean distance of
700 km from the coast. These long‐lived eddies can thus
impact the offshore ocean of the PCCS. On the basis of the
average values shown in Table 5, we estimate that the
annual volume anomaly of trapped fluid transported by
these 55 long‐lived eddies is of ∼7.5 Sv. Similarly, the
annual heat (salt, respectively) transport anomaly associated
with 30 long‐lived CEs would be of around −50 × 1011 W
(−90 × 103 kg s−1) and of 70 × 1011 W (190.0 × 103 kg s−1)
for 25 AEs. Note that since anomalies associated with CEs
and AEs are of opposite signs (the negative anomalies
indicate heat and salt deficiencies with respect to the background water mass), the total contribution of the mesoscale
eddies on the heat and salt transport is probably weak.
However, as clearly observed in Figure 8, heat and salt flux
anomalies are transported in distinct depth layers depending
on the eddy rotation. On average, mesoscale eddies would
contribute to inject relatively warm and salty water in the
subthermocline water (below 200–300 m depth) and to cool
and freshen the upper ocean above 200 m depth.
[34] The volume, heat and salt transports estimated for the
PCCS are relatively weak compared to more energetic
regions of the Southern Hemisphere such as in the Antarctic
Circumpolar Current [Joyce et al., 1981; Peterson et al.,
1982; Morrow et al., 2004; Swart et al., 2008], the Agulhas
Current [Gordon and Haxby, 1990; van Ballegooyen et al.,
1994; van Aken et al., 2003], the Brazil‐Malvinas region
[Gordon, 1989; Lentini et al., 2002; de Souza et al., 2006], or
in the South Atlantic subtropical gyre [McCartney and
Woodgate‐Jones, 1991]. However, our volume flux estimates are in agreement with the fluxes obtained in the similar
California Current System where model simulations have
Table 5. Thermohaline Contents and Associated Transports
Integrated Over the Volume of the Composite Cyclonic and
Anticyclonic Eddiesa
PCCS
(10°S–30°S)
CE
Vertical extent (m)
Volume (× 1012 m3)
AHA (× 1018 J)
ASA (× 1010 kg)
Volume transport (Sv)
Heat transport (× 1011 W)
Salt transport (× 103 kg s−1)
AE
Northern
Region
(10°S–20°S)
CE
AE
Southern
Region
(20°S–30°S)
CE
0–230 0–540 0–200 0–450 0–280
3.1
5.5
2.6
4.9
3.4
−5.5
8.7
−5.9
6.5
−5.3
−9.8 23.8 −14.7 17.4 −7.5
0.10 0.18 0.08 0.15 0.11
−1.7
2.8
−1.9
2.1
−1.7
−3.1
7.5
−4.7
5.5
−2.4
AE
0–570
6.1
10.5
28.3
0.19
3.3
9.0
a
Estimates were computed in the entire PCCS (10°S–30°S) and in the
northern and southern regions delimited by black heavy lines in Figure 1d.
AHA, available heat anomaly; ASA, available salt anomaly.
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Figure 9. Mean vertical profiles of (a, c) temperature anomaly and (b, d) salinity anomaly, inside CEs and
AEs as a function of their surface signature in dynamic height anomaly (relative to 1000 m). Anomalies are
projected either on depth (Figures 9a and 9b) or on s levels (Figures 9c and 9d). Grey lines in Figures 9a
and 9b correspond to s levels ranging from 24.8 to 27.2 kg m−3 with a contour interval of 0.2 kg m−3.
Black heavy lines correspond to the s levels delimiting the eddy cores where /S anomalies are maximum
(25.2–26.0 kg m−3 for CE and 26.0–26.8 kg m−3 for AE).
shown that the fluid trapped inside AEs extends from the
surface to 500 m depth and transport a volume flux of 0.25 Sv
[Cornuelle et al., 2000]. Note that our calculations only
represent part of the total eddy fluxes since we have not
included the possible large contribution of short‐lived eddies
to the volume, heat and salt budgets. Our results do indicate,
nonetheless, that mesoscale eddies can influence the zonal
heat and salt budgets of the PCCS.
5. Discussion
[35] In other eastern boundary current systems, and in
particular in the California Current System, it is now widely
accepted that nonlinear eddies are mainly generated by
instabilities of the near‐coastal currents [Batteen, 1997;
Huyer et al., 1998; Chereskin et al., 2000; Batteen et al.,
2003; Marchesiello et al., 2003; Capet et al., 2008; Kurian
et al., 2011]. In the California Current System, eddies generated by the surface California Current are predominantly
cyclonic with a surface core in the upper 150 m, whereas
those shed by the subsurface California Undercurrent are
mainly anticyclones with a subsurface core (400 m) [Simpson
and Lynn, 1990; Huyer et al., 1998; Garfield et al., 1999;
Chereskin et al., 2000; Cornuelle et al., 2000]. In the PCCS,
the instability of near‐coastal currents has been also considered as a key parameter for the generation of mesoscale
eddies that are mainly formed near the coast and propagate
westward [Leth and Middleton, 2004; Capet et al., 2008;
Johnson and McTaggart, 2010; Colas et al., 2011]. In fact,
two recent studies based on regional numerical simulations
have shown that modeled eddies in the California Current
System and the PCCS shared similar vertical characteristics
[Kurian et al., 2011; Colas et al., 2011]. In both these eastern
boundary current systems, modeled CEs were surface
intensified with their cores located in the upper 100–200 m
whereas AEs showed maximum subsurface /S anomalies
centered between 200 and 400 m depth [see Kurian et al.,
2011, Figures 16 and 17; Colas et al., 2011, Figures 14
and 15].
[36] In the present study, the composite analysis of Argo
CTD profiles corroborates the results obtained from regional
simulations, showing that in the PCCS the main core of the
CE is located above 200 m depth, whereas the core of the
AE is centered at 400 m depth (Figure 5). The vertical
position of the CE and AE cores is rather independent of the
eddy intensity, defined here as the surface DH′ value
(Figures 9a and 9b), and the maximum ′/S′ correspond to
density ranges of 25.2–26 kg m−3 for CEs and of 26–26.8 kg
m−3 for AEs (Figures 9c and 9d). The CE density class
(25.2–26 kg m−3) is associated with thermocline water
above the Eastern South Pacific Intermediate Water whereas
the AE density range (26.0–26.8 kg m−3) is related to subthermocline waters including both the Eastern South Pacific
Intermediate Water and the Equatorial Subsurface Water
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(Figure 1d). Once generated near the coast, one can assume
that the fluid trapped in the westward propagating eddies
transports /S anomalies mainly along isopycnal levels. But
near the coast, the density range associated with CEs (25.2–
26 kg m−3) corresponds to the relatively cold and fresh
upwelled water suggesting that intrathermocline CEs may
arise from the meandering of the Chile and Peru surface
currents separating Cold Coastal Water from the warmer and
saltier offshore water. The destabilization of these currents
can trap Cold Coastal Water from their inshore side, forming
CEs which can inject during their offshore displacements
relatively cold and fresh coastal water inside the thermocline/halocline. In contrast, near the coast, the density range
associated with the core of AEs (26–26.8 kg m−3) corresponds to the subsurface PCU flowing poleward [e.g.,
Johnson and McTaggart, 2010]. Thus, these subthermocline
AEs are more likely to be shed by the PCU [Colas et al.,
2011; Johnson and McTaggart, 2010].
[37] The eddy generation by the surface and subsurface
currents are also supported by three other particular features
observed in this study. First, in AEs maximum salinity
anomalies are found above the core of maximum temperature
anomalies (Figures 5–9). But in the offshore ocean, the
Eastern South Pacific Intermediate Water, corresponding to
the 26.0–26.3 kg m−3 s layer, is fresher and warmer that the
underlying Equatorial Subsurface Water associated with the
26.3–26.8 kg m−3 s layer (Figure 1d). Thus, an AE transporting a rather homogeneous water mass from the PCU
toward the offshore ocean, would produce, far from the coast,
a stronger salinity (temperature, respectively) anomaly in
the upper (lower) layer of density 26.0–26.3 kg m−3 (26.3–
26.8 kg m−3). Second, in agreement with a southward
deepening of the PCU core [Silva and Neshyba, 1979; Colas
et al., 2011], the composite AEs also show a deeper core in
the southern region (Figure 7). Third, the temperature and
salinity anomalies were stronger in the core of CEs (AEs,
respectively) in the northern (southern) region. From the
CARS climatology (not shown), we determined that the
temperature and salinity differences between the Cold
Coastal Water (PCU water, respectively) and the offshore
intrathermocline (subthermocline) are also stronger in the
northern (southern) region. Thus, the observed vertical
structure of the composite eddies confirms that CEs are most
likely formed by instabilities of the surface currents whereas
AEs are shed by the PCU.
6. Summary and Future Works
[38] On the basis of altimetry data and available Argo
profiles, this study has characterized the mean vertical
structure of both cyclonic and anticyclonic eddies in the
Peru‐Chile Current System. Through a composite analysis of
the Argo profiles located inside eddies we have highlighted a
key difference between the thermohaline vertical structures
of CEs and AEs: On average, the core of CEs is located in the
25.2–26.0 kg m−3 s layer corresponding to the thermocline,
whereas the core of AEs is found in the subthermocline
layer characterized by s of 26.0–26.8 kg m−3. Within their
core, these eddies exhibit typical temperature and salinity
anomalies of ±1°C and ±0.1 respectively, associated with a
geostrophic velocity on the order of ±10 cm s−1. Intrathermocline CEs are likely to be formed by instabilities of the
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near‐coastal surface currents and propagate offshoreward,
trapping in their cores recently upwelled cold coastal water.
In contrast and as also suggested by Johnson and McTaggart
[2010], subthermocline AEs seem to originate from the PCU
and to impact both the Eastern South Pacific Intermediate
Water and the Equatorial Subsurface water during their
propagation, with strongest salinity anomalies observed in
the layer of the relatively fresh Eastern South Pacific Intermediate Water and strongest temperature anomalies in the
underlying colder layer related to Equatorial Subsurface
Water. Spread over a 1 year period, the volume flux of near‐
coastal water trapped within each eddy is of 0.1–0.2 Sv. The
heat and salt transport anomalies associated with each of
these typical eddies are about ±2–3 × 1011 W and ±3–8 ×
103 kg s−1 depending on their rotation. Negative anomalies
are rather found in the upper thermocline level whereas
positive anomalies are located in subthermocline levels.
[39] The results may be useful for the validation of high‐
resolution regional models [e.g., Colas et al., 2011] and
could be of interest to the biogeochemical community to
investigate links between ecosystems and mesoscale eddies
in the highly productive Peru‐Chile Current System. In turn,
regional models could help to document the exact mechanisms involved in the formation of both the CEs and AEs in
the near‐coastal region. These simulations could also be used
to more precisely estimate the cross‐shore transports of heat
and salt associated with both the mesoscale eddies and the
large‐scale circulation.
[40] The results also raise a number of additional questions that require further investigation. For example, the core
of AEs is located in the Equatorial Subsurface Water layer
associated with the shallowest and most pronounced oxygen
minimum zone of the world ocean [Helly and Levin, 2004;
Fuenzalida et al., 2009; Stramma et al., 2010]. This hypoxic
layer influences biogeochemical cycling of elements and has
a strong impact not only on the local rich ecosystem but also
on the global climate, being a significant source of active
greenhouse gases (CO2 and N2O) [Paulmier and Ruiz‐Pino,
2009]. One can thus expect that AEs also impact the oxygen‐
minimum zone through the offshore propagation of dissolved oxygen anomalies from the near‐coastal region.
Similarly, CEs could also impact the biogeochemical tracer
budgets through the offshore advection of relatively nutrient‐
rich and oxygen‐poor cold coastal water into the thermocline of the offshore ocean. Thus, the impact of both CEs
and AEs on the nutrients and dissolved oxygen contents
should also be addressed in future research. These studies
should in particular focus on the decay of mesoscale eddies,
which is a key process for the redistribution of trapped
properties toward the surrounding water [e.g., Whitney and
Robert, 2002; Johnson et al., 2005; Swart et al., 2008;
van Sebille et al., 2010; Early et al., 2011].
[41] Acknowledgments. Float data used here were collected and
made freely available by Argo (http://www.argo.net/), a program of the
Global Ocean Observing System, and contributing national programs.
The altimeter products were produced by Ssalto/Duacs and distributed by
AVISO, with support from CNES. The preparation and development of this
work were made possible thanks to stays of A. Chaigneau and M. Le Texier
at IMARPE. Several Argo float deployments have been supported by
GMMC (Mercator‐Coriolis) through the Flotteurs du Pacifique Sud
(FLOPS) project. O.P. was supported by FONDECYT 1090791. Support
from ECOS‐CONICYT is also acknowledged. This work was partially
14 of 16
C11025
CHAIGNEAU ET AL.: EDDY VERTICAL STRUCTURE IN THE ESP
funded by the LMI “Dinamica del Sistema de la Corriente de Humboldt.”
The authors particularly thank F. Colas, D. Chelton, and E. van Sebille for
their constructive comments that considerably helped to improve a previous
version of the manuscript. We are also grateful to R. Samelson and
D. Chelton for interesting discussions on the definition of the “eddy trapping
depth.”
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