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Equatorial Asymmetry of the East Pacific ITCZ: Observational Constraints
on the Underlying Processes
HIROHIKO MASUNAGA
Hydrospheric Atmospheric Research Center, Nagoya University, Nagoya, Japan

TRISTAN S. L’ECUYER
Department of Atmospheric Science, Colorado State University, Fort Collins, Colorado(Manuscript received 17 May 2010, in final form 13 November 2010)
ABSTRACTThe equatorial asymmetry of the east Pacific intertropical convergence zone (ITCZ) is explored on the basisof an ocean surface heat budget analysis carried out with a variety of satellite data products. The annual mean
climatology of absorbed shortwave flux exhibits a pronounced meridional asymmetry due to a reduction of
insolation by high clouds in the north ITCZ. Ocean mixed layer advection has the largest, if not exclusive,
effect of counteracting this shortwave-exerted asymmetry. Other heat fluxes, in particular latent heat flux,
predominate over the advective heat flux in magnitude but are secondary with respect to equatorial asymmetry. The asymmetry in advective heat flux stems from a warm pool off the Central American coast and, toa lesser extent, the North Equatorial Counter Current, neither of which exist in the Southern Hemisphere.
The irregular continental geography presumably comes into play by generating a warm pool north of theequator and bringing cold waters to the south in the far eastern Pacific.
In addition to the annual climatology, the north–south contrast in the seasonal cycle of surface heat flux is
instrumental in sustaining the north ITCZ throughout the year. The northeast Pacific is exposed to a seasonal
cycle that is considerably weaker than that in the southeast Pacific, arising from multiple causes including thefinite eccentricity of the earth’s orbit and meridional gradient in mixed layer absorptivity. Simple experiments
generating synthetic sea surface temperature (SST) illustrate that the muted seasonal cycle of heat flux forcing
moderates the SST seasonal variability in the northeast Pacific and thus allows the north ITCZ to persist year
round. Existing theories on the ITCZ asymmetry are briefly examined in light of the present findings.

1. Introduction
The equatorial asymmetry of the intertropical convergence zone (ITCZ) over the east Pacific Ocean has been
extensively studied over past decades. A tight correlationof atmospheric dynamics with underlying sea surface temperature (SST), as well known from observations (Fig. 1),
offer compelling evidence for the critical importance of
air–sea interaction. Different hypotheses have been proposed to explain the ITCZ asymmetry in terms of air–sea
interaction including the wind–evaporation–SST (WES)
feedback (Xie and Philander 1994), stratus–SST feedback (Philander et al. 1996), and upwelling–SST feedback

Corresponding author address: Hirohiko Masunaga, Hydrospheric Atmospheric Research Center, Nagoya University, F31(200) Furocho Chikusa-ku, Nagoya 458-0015, Japan.
E-mail: masunaga@hyarc.nagoya-u.ac.jp
DOI: 10.1175/2010JCLI3854.1
Ó 2011 American Meteorological Society(Mitchell and Wallace 1992; Chang and Philander 1994),
as reviewed by Li (1997) and Xie (2004). The irregular continental geometry along the Central and SouthAmerican coastlines arguably destroys the equatorial
symmetry in addition to the processes internal to ocean–
atmosphere dynamics (Philander et al. 1996; Xie andSaito 2001).
The ITCZ asymmetry has been studied principally bymeans of coupled ocean–atmosphere models at different
levels of complexity. Simplified models are suited for
efficiently tracking down critical physical processes, but
the simplifications could filter out some processes that
otherwise might be found crucial for giving rise to theasymmetry. Meanwhile, it remains a formidable task forstate-of-the-art coupled climate models to simulate theeast Pacific ITCZ precisely as observed (Mechoso et al.
1995; Dai 2006; de Szoeke and Xie 2008). Comparedto this rich spectrum of modeling studies exploring the15 MARCH 2011

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in situ stations. Such an approach is required since the
ITCZ asymmetry essentially involves large-scale air–sea
dynamics.
In this work, the observed equatorial asymmetry of
the east pacific ITCZ is documented in terms of theestimated heat fluxes and underlying factors such as
clouds, wind, and SST. Following an overview of theanalysis strategy in section 2, the annual mean climatology is first presented (section 3) to outline theprocesses that potentially maintain the north–south
asymmetry of the ITCZ. Factors responsible for the
ITCZ asymmetry in the annual mean field are further
examined in comparison with the seasonal double ITCZ
observed in the March–April climatology (section 4).
The persistence of the north ITCZ and the transienceof the south ITCZ are then explored by analyzing theseasonal cycle of the surface heat fluxes in section 5.
The present findings are discussed and concluded in
section 6.
FIG. 1. (a) Geographical map of annual-mean deep storm coverage (shaded) and SST (contour). The contour interval is 18C, and278C is highlighted by the thick line. The dashed box indicates theprimary target area of the present analysis. (b) As in (a) but for
March–April bimonthly mean.

ITCZ asymmetry, observational work against which proposed hypotheses can be tested in depth is rather limited.
A notable exception is the East Pacific Investigation of
Climate Processes in the Coupled Ocean–Atmosphere
System 2001 (EPIC2001), a field campaign program
targeted on interactive air–sea processes in the tropicaleast Pacific (Raymond et al. 2004). The present paper isintended to join efforts to fill in the lack of such observational benchmarks as well as to identify key processes possibly missing from the existing theories. To
this end, the ocean mixed layer heat budget is analyzed
with surface heat fluxes estimated primarily from a wide
variety of satellite measurements.
The ocean surface heat budget has been studied extensively by combined in situ and satellite measurements
over the equatorial Pacific (Wang and McPhaden 1999),
the equatorial cold tongue (Swenson and Hansen 1999),
the ITCZ–equatorial cold tongue complex (McPhaden
et al. 2008), and the east Pacific warm pool (Wijesekera
et al. 2005), among others. The current analysis relies onsatellite data and ancillary global datasets and therefore
does not attempt to quantitatively refine the previous
heat budget estimates derived from buoy and ship observations. This work, instead, is aimed at analyzing theclimatological large-scale structure of surface heat fluxes
and ocean heat transport over the entire tropical east
Pacific beyond the limited areal coverage observable by

2. Data and method
a. Ocean heat flux estimate
The ocean mixed layer heat budget is analyzed witha variety of datasets based primarily on satellite measurements. The overall strategy follows Masunaga and
L’Ecuyer (2010) but a number of modifications have been
made to improve the fidelity of the model. The basic
equation governing the mixed layer heat budget isrw H ml(x, y)Cp,w›T s(x, y, t)
5 Qnet(x, y, t)  hQres i(x, y).
›t(1)

The net heat flux, Qnet, is broken down into
Qnet 5 Qadv 1 Qupw 1 QSW 1 QLW 1 QLH 1 QSH 1 Qdif ,where Qadv is the advective heat flux, Qupw the upwellingheat flux, QSW the shortwave flux absorbed into the ocean
mixed layer, QLW the longwave flux, QLH the latent heat
flux, QSH the sensible heat flux, and Qdif the diffusional
heat flux. The residual term,
hQres i(x, y) 5þ1
Qnet (x, y, t) dt,
T

represents biases associated with the present estimates ofindividual heat fluxes. Here T represents one climatological year and the integral applies to the whole annual
cycle. Methodology to derive individual heat fluxes is next
reviewed.

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1) LATENT AND SENSIBLE HEAT FLUXESThe latent heat flux through evaporation from theocean surface is estimated by use of a bulk formula:)]ju10 j ,QLH 5 ra Ly CE [qy  q*(T
y
s

(2)

where ra, Ly, qy, q*y (Ts), and u10 are the dry air density,
vapor latent heat per unit mass, vapor mixing ratio, saturation vapor mixing ratio for SST or Ts, and wind velocity
at 10 m above the sea surface, respectively. The bulk
transfer coefficient for evaporation, CE, is adopted from
Large et al. (1994). In this study, Ts and u10 are provided
by the Tropical Rainfall Measuring Mission (TRMM)
Microwave Imager (TMI) and the Quick Scatterometer(QuikSCAT) SeaWinds instrument, respectively, andqy is related to TMI-retrived column water vapor (CWV)
through the water vapor scale height, Hy, asra qy 5 CWV/H y ,(3)where Hy is assumed to be a function of CWV and SST:
H y 5 f (CWV, T s ).

(4)

Masunaga and L’Ecuyer (2010) found that Hy is wellconstrained by CWV and SST, and empirically defineda bivariate linear regression equation for (4) in comparison with moored buoy data. This linear formula isreplaced in the present work by a tabulated database
representing f(CWV, Ts), constructed with the Atmospheric Infrared Sounder (AIRS) daily gridded productmatched to CWV and SST from the Advanced Microwave Scanning Radiometer for the Earth Observing
System (AMSR-E). The AIRS vapor mixing ratio in
the lowermost atmospheric layer (1000–925 hPa) is extrapolated down to the surface using a simple boundary
layer model, in which potential temperature and vapor
mixing ratio are assumed to be constant within the atmospheric mixed layer and follow a moist adiabat above.
The surface vapor mixing ratio, providing the lower
boundary condition to this model, is evaluated so thatthe model temperature and humidity profiles are consistent with a given AIRS estimate. Simultaneous AIRS andAMSR-E measurements over the tropical east Pacific forfive years from 1 January 2003 to 31 December 2007 are
employed to build the lookup table of Hy.
The sensible heat flux, QSH, is taken from the Japanese
Ocean Flux Datasets with Use of the Remote Sensing
Observations 2 (J-OFURO2) dataset (Kubota et al. 2002).

where Cp,w 5 4.0 3 103 J kg21 K21 is the specific heat ofocean water, rw 5 1030 kg m23 the ocean water density,u the current velocity taken from the Ocean Surface
Current Analysis–Real time (OSCAR) product (Bonjean
and Lagerloef 2002), and Hml is the mixed layer depth
adopted from de Boyer Montégut et al. (2004).
The upwelling velocity, wE, at the mixed layer bottom is
wE 5

Qadv 5 Cp,w rw H ml u  $T s ,1($  Mml ),rw(6)

where the mixed layer mass transport, Mml, is determined
as follows. The conventional formulation of 1.5-layer
models for tropical oceans implies
um 5 u1  ue

H tc  H ml
,
H tc

(7)

where um is the vertical mean current in the dynamically
active layer, u1 is the surface current, and ue is the Ekman
current or the shear between the mixed layer and underlying layer (e.g., Zebiak and Cane 1987; Chang 1994).
The thermocline depth, Htc, is approximated by the 208C
depth determined with the World Ocean Atlas 2005
(WOA05) (Locarnini et al. 2006). Equation (7) is rewritten in terms of mass transport as
H ml
H  H ml
M 5 Mml  tc
ME ,
H tc
H tc

(8)

where M [ Htcum is the total mass transport above thethermocline and ME [ Hmlue is the surface Ekman masstransport. Taking the divergence of (8), one obtainsH ml
H  H ml, (9)$  M 5 $  Mml  $  ME tc
H tc
H tc
where Hml/Htc is assumed not to be highly variable overdistance. In (11), the lhs vanishes if $  M is negligibly
small compared to individual terms on the rhs, that is, theactive layer is nearly nondivergent as a whole. Combining(6) and (9) then leads toH  H ml1.
(10)
$  ME tc
wE 5
rw
H tc
The surface Ekman flow, ME [ (ME,x, ME,y)T, isevaluated from the observed wind stress field as2) OCEAN ADVECTION AND UPWELLING FLUXES
The advective heat flux is evaluated as

VOLUME 24

r s M E, x  f M E, y 5 t x(11)

rs ME, y 1 f ME, x 5 t y ,

(12)

and
(5)

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where r s 5 6.62 3 10 26 s 21 is the drag coefficient(Lagerloef et al. 1999), f is the Coriolis parameter, and t x
and t y are the zonal and meridional components of thewind stress, respectively. The bulk formula by Large et al.
(1994) is applied to the QuikSCAT wind vector to estimate t x and t y in (11) and (12). The upwelling (or entrainment) heat flux is then evaluated from wE as
Qupw 5 rw Cp,w W e H(W e )DT E ,W e 5 wE 1 u  $H ml ,where H is the Heaviside function and DTE represents thetemperature difference between mixed layer water andupwelling water from beneath. In the entrainment rate,
the Eulerian derivative term ›Hml/›t has been omitted
since Hml is temporally fixed in the current analysis. The
second term in We is minor, overall, compared to the
upwelling velocity. In this work, DTE is defined between
the depths of 0.5Hml and Hml 1 15 m using the WOA05
temperature profile.

3) RADIATIVE FLUXThe downward shortwave and longwave fluxes at thebottom of the atmosphere (BOA), denoted F YSW and F YLW
are provided by the Hydrologic Cycle and Earth Radiation Budget (HERB) product (L’Ecuyer and Stephens
2003, 2007; L’Ecuyer and McGarragh 2010). The net
longwave heat flux at the surface is written as
QLW 5 (1  s )F YLW  s sT 4s ,(14)

where s 5 5.67 3 1028 W m22 K24 is the Stephan–Boltzmann constant and s is the longwave surface emissivity. In this study, s is fixed at unity (Kraus and Businger
1994; Large and Yeager 2009). The net shortwave flux
absorbed by the ocean mixed layer is
QSW 5 (1  a)PF YSW ,where g 5 0.47 is the fraction of solar flux that falls in thepenetrative wave band (380–700 nm) and hg 5 20 m isthe depth scale for the penetration of insolation over thiswave band (Murtugudde et al. 2002).

4) DIFFUSIONAL HEAT FLUXESThe ocean heat flux due to horizontal diffusion is(13)

with the entrainment rate at the mixed layer base
(Stevenson and Niiler 1983):(15)

where a is the ocean surface albedo and P is the fractionof incoming shortwave energy absorbed into the mixed
layer. Ocean surface albedo is specified as a function ofsolar zenith angle using the spectral albedo model of Jin
et al. (2004). For simplicity, P is parameterized by a simple function of Hml:
"

!#
!
H ml
H ml
5 1  g exp ,
P 5 (1  g) 1 g 1  exp 
hg
hg
(16)

1787

Qdif 5 rw Cp,w H ml k=2 T s ,(17)where the turbulent diffusion coefficient k is conservatively chosen to be 3 3 103 m2 s21 (e.g., Bauer et al.
1998). Vertical diffusion at the mixed layer base is notconsidered in the present analysis.

b. Data
TMI-retrieved SST and CWV datasets are provided by
the Remote Sensing Systems (RSS), and u10 is from the
QuikSCAT L3 daily data product. Low and high cloud
fractions analyzed in this paper are retrieved from TRMM
TMI and Visible/Infrared Scanner (VIRS) measurements
by the HERB algorithm and are, therefore, physically consistent with the radiative flux estimates. Precipitating portions of cloud systems are delineated by use of the TRMM
Precipitation Radar (PR) echo-top height along with VIRS
infrared brightness temperature (IR Tb) (Masunaga et al.
2005; Masunaga and Kummerow 2006). Vertically developed precipitating clouds, represented by the combinationof the ‘‘deep stratiform’’ and ‘‘deep convective’’ classes in
the Masunaga et al. (2005) terminology, are identified by
PR echo tops higher than 4 km collocated with IR Tb colder
than 245 K. The area fraction of such towering precipitating clouds is hereafter denoted by deep storm coverage and used as an indicator of deep convective activity.
Note that the deep storm coverage only represents precipitating clouds while the HERB cloud fractions contain
nonprecipitating clouds as well.
The whole dataset is analyzed on a 18 3 18 Cartesian
grid for eight years of observations from 1 January 2000
to 31 December 2007. The J-OFURO2 dataset is unavailable beyond 2006, so the sensible heat climatology iscomputed from seven years omitting 2007. All heat fluxes
and atmospheric parameters studied are first evaluatedon a daily basis and then averaged to climatology. Pentad
time series of the OSCAR current is temporally interpolated to daily sequences. Ancillary parameters for characterizing the ocean interior, that is, Hml, DTE, and Htc, arefixed at the annual mean field.
The primary study area is zonally bound by 1258 and
1008W, chosen so that the analysis is targeted specificallyon the east Pacific ITCZ. The south ITCZ, when it appears,
would be contaminated with the easternmost portion of the

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FIG. 2. Climatological zonal-mean heat fluxes averaged between 1258 and 1008W. Despite having different intercepts, the vertical axis
spans 200 W m22 in all panels. (a) Total BOA shortwave flux (dashed) and that absorbed at the surface (QSW, solid). (b) Latent heat flux
(QLH, solid), upwelling heat flux (Qupw, dotted), longwave flux (QLW, dashed), and advective heat flux (Qadv, dotted–dashed). (c) Sensible
heat flux (QSH, solid), diffusion heat flux (Qdif, dashed), and the residual term (dotted–dashed).

South Pacific convergence zone (SPCZ) if the westernedge were extended beyond 1258W, while the easternedge (1008W) of the study area is limited by a geographical constraint imposed by the Central American landmass (see Fig. 1). The sectors south and north of theequator dividing the study area are hereafter denoted by
the southeast Pacific (SEP) and the northeast Pacific(NEP). In section 5, the SEP (NEP) is limited further tothe meridional range of 158S–28S (28N–158S) for defining
regional means, excluding the narrow equatorial band
between 28S and 28N to avoid the equatorial cold tongue.

3. Annual mean climatology
The annual mean climatology of the east Pacific ITCZ
is first examined focusing on the different factors contributing to the ocean heat budget.

a. Zonal-mean heat fluxes
Figure 2 shows the zonal-mean heat fluxes averaged
from 1258 to 1008W. The incoming shortwave flux at the
BOA (dashed line in Fig. 2a) exhibits a remarkable meridional asymmetry, showing a distinct minimum around
58–108N and a maximum at the equator. This equatorial
asymmetry is preserved as the surface shortwave flux
penetrates deeper and is absorbed in the ocean mixed
layer, reaching a maximum difference of ;40 W m22
between 108S and 108N (solid line in Fig. 2a). Since insolation at the top of the atmosphere is virtually symmetric
about the equator when the seasonal cycle is averaged out,
this north–south asymmetry results from unevenly distributed atmospheric and oceanic components that modulate the downwelling shortwave flux absorbed in themixed layer, as will become clearer later.
The energy input by the absorbed shortwave flux must
be locally balanced in the long-term mean with all otherheat fluxes combined. Figure 2b illustrates major heat

fluxes counteracting the absorbed shortwave flux. The
latent heat flux is predominant (i.e., largest in negativeflux) except near the equator, where the upwelling cooling
dominates the other terms. The latent heat flux is characterized by a pair of negative maxima near 108S and 108N
with the southern maximum slightly exceeding its northern counterpart in negative magnitude. The longwave flux
is relatively homogeneous across the study area. The advective heat flux, although minor in amplitude comparedto the other fluxes plotted together, has a notable north–south contrast and even changes sign to positive for latitudes north of 78N.
The remaining heat fluxes and residual term are depicted in Fig. 2c. The sensible and diffusion heat fluxes areminimal throughout the entire tropical latitudes. The residual term, hQresi, exhibits an equatorially asymmetric
twist in proximity of the equator (58S–58N) but transitions
into a roughly symmetric pattern as latitude increases
farther poleward . The equatorial bias is likely attributed toerrors in the upwelling flux estimate, subject to our limited
knowledge of rs, DTE, Hml, and Htc, and possibly to thevertical diffusional flux as well, which is ignored in thispaper. A number of experiments have been performed
with tunable parameters such as P, rs, and DTE perturbed
within a plausible range (not shown). The residual termchanges magnitude and even in sign, when those parameters are varied one after another, but the analysis results
critical to the major conclusions of this study are qualitatively insensitive to the perturbations imposed. The current findings would therefore be similar if the same
analysis were repeated with an ideal dataset subject to no
closure error. The shortwave and latent heat fluxes, having
a predominant magnitude over other fluxes, are perhaps
the main sources of the bias poleward of 108S and 108N.
The residual term becomes nonnegligibly large near 208S
and 208N, while our present interest is limited to latitudes
lower than 158 where the ITCZ mostly stays.

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FIG. 3. North–south difference in climatological zonal-mean(1258–1008W) heat fluxes computed for each latitude denoted by
the abscissa. North exceeds south when the differential flux is
positive. Plotted are absorbed shortwave flux (QSW, shaded histogram), latent heat flux (QLH, solid), upwelling heat flux (Qupw,
dotted), longwave flux (QLW, dashed), and advective heat flux
(Qadv, dotted–dashed).

The equatorial asymmetry is quantitatively demonstrated
in Fig. 3, where the difference between the Northern andSouthern Hemispheres is computed at each latitude forindividual heat fluxes. The absorbed shortwave flux
(shaded) has a negative asymmetry (i.e., the SEP flux
exceeds the NEP flux) with the magnitude largest among
all of the heat fluxes. In contrast, the advective, longwave,
and latent heat fluxes show a positive asymmetry between
58 and 158 in latitude. Although it is least significant whencompared in absolute value (Fig. 2b), the advective heat
flux has the largest equatorial asymmetry among thesethree fluxes and is therefore the most important heatbudget component that counteracts the asymmetry exerted by the absorbed shortwave flux. The upwelling heat
flux dominates the narrow equatorial area at latitudes
lower than 58, beyond which the upwelling cooling rapidly tapers off. Given that the latitude at 58 meridionallyseparates the ITCZ from the equatorial cold tongue, it issuggested that, while the advective heat flux may play a
crucial role in the ITCZ asymmetry, the contribution of
upwelling cooling is limited to the equatorial cold tongue.

b. Factors controlling the shortwave modulation
The meridional modulation of absorbed shortwave flux
is ascribable to multiple factors including high and low
cloud fractions and mixed layer absorptivity. Figures 4a–c
show the climatological maps of total cloudiness and its
decomposition into high and low cloud fractions. The
geographical distribution of the high cloud fraction is1789

clearly delineated in the northern ITCZ region, while thelow cloud fraction is enhanced over cold oceans wheremaritime stratocumuli often develop. Note that the HERB
product uses passive microwave and visible–infrared measurements and therefore identifies both high and low
clouds even where they coexist.
The large-scale variability of the BOA shortwave flux
(Fig. 4d) is apparently explained by the shielding effects
of high clouds extending along the north ITCZ as well as
stratocumulus cloud decks. The spatial pattern of theshortwave flux is slightly modulated when multiplied by
the mixed layer absorptivity (Fig. 4e), leading to a further
reduction of the absorbed shortwave flux around 58–108N,
as seen in Fig. 2a, due to a systematic northward decreaseof the mixed layer absorptivity [Fig. 4f, see also Fig. A2 in
Masunaga and L’Ecuyer (2010) for the spatial distribution of the mixed layer depth].
Although the conspicuous shortwave reduction in the
ITCZ has been long recognized (Oberhuber 1988; Mitchelland Wallace 1992; Cronin et al. 2006), high clouds extending along the ITCZ are often considered more asa result of the ITCZ asymmetry and less as a cause forit. The ITCZ clouds, however, appear to play a key rolein the shortwave modulation for the study area of thepresent work. Since high clouds in the ITCZ are mainly
associated with deep convection, their radiative impact
is not just an external forcing to the mixed layer heatbudget but relies on the preexisting ITCZ, suggesting
the presence of feedback processes that maintain theobserved equatorial asymmetry.

c. Factors controlling the latent heat flux
The latent heat flux given by (2) may be decomposed
into the scalar wind speed, ju10j, and the air–sea humidity
difference, [q*y(Ts) 2 qy], as principal parameters responsible for its variability. These two components are separately plotted in Fig. 5. The zonal-mean wind speed andhumidity difference are both double peaked, straddling anequatorial minimum and paralleling the meridional
structure of the latent heat flux itself (Fig. 2b). The spatial
variability of the air–sea humidity difference (Fig. 5b) is
controlled mainly by the saturation vapor term q*y (Ts), as
inferred from the resemblance to the SST distribution(Fig. 1). The wind field in the study area consists of thesouthern and northern trades, which eventually converge
into the ITCZ (Fig. 5b). The southeasterly, cross-equatorial
trade decelerates as it approaches the equator, giving
rise to the aforementioned wind minimum, before accelerating again beyond the equator (Wallace et al. 1989;
Liu and Xie 2002).
Despite their mutual resemblance in the meridional
pattern, the zonal-mean wind speed and humidity difference exhibit a subtle equatorial asymmetry in opposing

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FIG. 4. Climatologies of (a) total cloudiness, (b) high cloud fraction, (c) low cloud fraction, (d) BOA shortwave flux (W m22), (e)absorbed shortwave flux (W m22), and (f) mixed layer absorptivity (P). A darker shade corresponds to a smaller shortwave flux in (d) and(e), so shortwave flux is better compared with cloudiness in (a)–(c). The study area is indicated by the dashed box.

ways. The northern peak is slightly stronger than thesouthern peak in the air–sea humidity difference, while
this asymmetry is reversed for the wind speed. Both effects are likely to arise from the presence of the north
ITCZ where SST is higher (and thus the air–sea humidity
difference larger) and the surface wind is weaker than on
the opposite side of the equator. When combined into thelatent heat flux, the wind speed asymmetry dominates that
due to air–sea humidity.

The dynamics governing the surface wind over thetropical east Pacific was discussed by Xie and Philander(1994) in the context of explaining the ITCZ asymmetry.
In a hypothesis later called the WES feedback, a sea levelpressure gradient induced by the north ITCZ arguably
enhances the wind speed in the SEP as the result ofa southerly wind with a westward veer superposed on the
background easterly trade. Observational evidence forthe wind–ITCZ correlation is sought in month-to-month

FIG. 5. (a) Climatological zonal-mean (1258–1008W) scalar wind speed (m s21) (ju10j, solid) and air–sea humidity
difference (g m23) (ra(q*(Tys) 2 qy ), dashed). Note that the humidity difference is defined with the dry-air density, ra,multiplied. (b) Climatological 10-m wind vector and air–sea humidity difference contoured every 0.5 g m23. The
study area for (a) is indicated by the dashed box. It should be noted that the climatological wind speed drawn in (a) is
not jhu10ij but hju10ji, where the bracket designates the temporal mean, so that it does not necessarily agree with thenorm of the wind vectors in (b).

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1791

FIG. 6. (a) The time series of monthly mean deep storm coverage (%) for the north ITCZ region (28–158N and
1258–1008W). Red circles and blue squares indicate active months and inactive months, respectively, as defined bydeviation from the mean with the one-sigma threshold. (b) Zonal-mean (1258–1008W) wind speed (shaded histogram,labeled on left axis) and latent heat flux (lines, labeled on right, upside-down) anomalies for active months (yellow/red) and inactive months (blue).

variability shown in Fig. 6. The time series of deep storm
coverage averaged over the NEP, as a proxy of the north
ITCZ activity, exhibits a pronounced temporal (primarily
seasonal) variation. Active (inactive) months can be defined as periods when deep storm coverage is larger(smaller) than the all-month mean (2.64) by at least one
standard deviation (1.32). The zonal-mean wind speed
and latent heat flux averaged for active/inactive months
are plotted as anomalies from climatology in Fig. 6. It isclear that wind speed is enhanced in the SEP when thenorth ITCZ is active, while this pattern is almost precisely
reversed when inactive. The latent heat flux closely follows this variation, deviating from the climatology by 215
or 120 W m22 near 108S. Although correlation does notnecessarily imply causality, the wind speed probably responds to the ITCZ activity on a monthly mean basis. This
result per se, however, does not guarantee that the WES
feedback is entirely responsible for the ITCZ asymmetry,as will be discussed further in section 6.

of days sampled, add up to the total (unconditional) mean(i.e., Fig. 7a) when summed together. The easterly conditioned composite (Fig. 8a) shows a broad patch of positive advective heat flux over the NEP, ascribable to theNorth Equatorial Current (NEC) carrying warm waterfrom an area of high SST offshore from Central America,or the eastern Pacific warm pool. When conditioned by
the westerly current, on the other hand, the advective
heat flux is more neutral but makes a small contribution
to the positive heat transport north of 58N, fed by thed. Factors controlling the advective heat flux
Figure 7a shows, as found in Fig. 2b, that the advective
heat flux has a notable meridional asymmetry characterized by a positive heat flux in the ITCZ and to its north.
To isolate the origins of this pattern, the advective heat
flux is broken down into SST and current velocity in Fig.
7b. A negative advective heat transport dominating the
SEP and equatorial band is obviously brought by theSouth Equatorial Current carrying cold water to the west
from near the South American coast and equatorial cold
tongue. An area of positive heat transport north of 78N issomewhat more complicated to interpret. In an attempt
to clarify its origins, the easterly and westerly conditioned
composites of advective heat flux and current velocity are
individually plotted in Fig. 8, where days with a negative(positive) zonal current are isolated to construct theeasterly (westerly) conditioned composite. Note that theadvective flux composites, each weighted by the number

FIG. 7. Climatologies of (a) advective heat flux (W m22) and (b)SST (8C) (colored) overlaid with ocean current velocity (cm s21).
The study area is indicated by the dashed box.

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FIG. 8. Composite advective heat flux (W m22) and current velocity (cm s21) conditioned by(a) easterly and (b) westerly currents. See text for the procedure of the conditional averages.
The study area is indicated by the dashed box.

North Equatorial Counter Current (NECC) that flows
along the ITCZ. As such, the warm water advection in thenortheast Pacific comprises two separate sources: one
originating in the east Pacific warm pool through the
NEC and, to a lesser extent, another linked with the west/
central Pacific ITCZ through the NECC.
We are now left with two questions to address in orderto understand the root of the meridional asymmetry in theadvective heat flux. The first is why the east Pacific warm
pool exists only in the NEP, and the other is why the SEP
lacks a band of westerly current as the NECC. The firstquestion requires an explanation for the mechanisms
maintaining the east Pacific warm pool. The meteorology of the wind specific to this area is of critical importance among processes governing the SST (Xie et al.
2005), and therefore the ocean heat budget there is sensitive to topographical constraints including continental
geography and local orography that directly dictate thewind and coastal current. The geographical configurationof the landmass is highly asymmetric at the eastern endof the tropical Pacific, and has been found by coupledmodel experiments to impose a strong constraint on the
ITCZ asymmetry (Philander et al. 1996; Xie and Saito
2001). The second question is rather a classical problem
belonging to Sverdrup’s theory. The equatorial doldrums
in the ITCZ are ultimately responsible for the NECC
underneath as a result of a negative (positive) wind stresscurl induced to its south (north). It follows that the absence of a southern counterpart of the NECC is attributed to the fact that the SEP has no permanent ITCZ.

4. Austral fall climatology and double ITCZ
Potential causes for the equatorial asymmetry in individual heat fluxes were discussed in some detail in theprevious section. High cloudiness associated with deep
convection has the greatest influence on the shortwave
modulation in the ITCZ. The surface wind is the principal element controlling the latent heat flux and shows amonth-to-month variation closely related to the ITCZ
activity. The meridional asymmetry in advective heat flux
is sustained by a heat transport by the NEC and anotherby the NECC. Whereas geographic origins presumably
account for the fact that the NEC is fed with warm waterand the SEC with cold, all other factors mentioned above
require a preexisting asymmetry of the ITCZ. If the ITCZ
were on both sides of the equator in the first place, one
might expect that the resultant fields of high cloudiness,wind speed, and equatorial counter current would besymmetric as well, being in balance rather than driving
toward the asymmetric state observed. Indeed, Fig. 6
implies that the wind reduction in the SEP for the monthsof an inactive ITCZ is large enough to offset the climatological equatorial asymmetry of latent heat flux.

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FIG. 9. (a),(b) As in Figs. 2a,b but for the bimonthly climatology for March and April. (c) As in Fig. 3 but for the bimonthly climatology
for March and April.

This hypothetical symmetric condition may be exemplified by the seasonal double ITCZ that appears in austral fall, as seen in Fig. 1b (Masunaga and L’Ecuyer 2010,
and references therein). Figure 9 shows the March–Aprilbimonthly climatology of zonal-mean heat fluxes as inFig. 2. The BOA shortwave flux exhibits paired minima
symmetrically aligned at 58S and 58N, according to thelatitudinal location of the ITCZ in this season. Thissymmetry is slightly skewed when absorbed into theocean mixed layer as a result of the north–south gradientof mixed layer absorptivity. The latent heat flux is also
symmetric about the equator within latitudes containing
the ITCZ (;108S–108N). In contrast, the advective heat
flux retains some asymmetry, as seen in the annual mean
field, with a distinct positive peak associated only with thenorth ITCZ. Figure 10a presents the advective heat flux
averaged for March and April. Unlike the annual mean
field shown in Fig. 7a, a positive heat transport is evident
off the South American coast, arising from the SEP warmband observed only in this season (Fig. 10b, see also
Masunaga and L’Ecuyer 2010). This warm water transport, however, is not as efficient as is in the NEP, presumably owing to the transience of the south ITCZ as wellas to the absence of a prominent warm pool in the SEP.
There is also no recognizable South Equatorial Counter
Current (SECC) despite the presence of a double ITCZ at
this time of year. If a SECC existed, it could supply thesouth ITCZ with warm water across a large SST gradientin the SPCZ and enhance the advective heat flux in the
SEP, making its zonal mean profile more symmetric
about the equator. This appears not to be the case because the south ITCZ does not persist long enough for
ocean dynamics to attain Sverdrup equilibrium in response to the rapidly changing atmospheric forcing.

and disappearance of the south ITCZ are primarily controlled by the seasonal cycle of insolation. An outstanding
question concerning the equatorial asymmetry of the
ITCZ is what keeps the north ITCZ active throughout
the year against the seasonally varying solar forcing. The
seasonal cycle of heat flux forcing is examined next toaddress this problem.

a. Seasonal cycle
Figure 11 shows the climatological annual time seriesof ocean surface heat fluxes averaged over the SEP and
NEP individually. The residual in the annual mean flux

5. On the persistence of the north ITCZ
Masunaga and L’Ecuyer (2010) showed observational evidence demonstrating that the development

FIG. 10. As in Fig. 7 but for the bimonthly climatology for Marchand April.

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FIG. 11. Climatological annual time series of (a) net heat flux (with the residual term subtracted), (b) absorbed
shortwave flux, (c) advective heat flux, and (d) upwelling, longwave, and latent heat fluxes as indicated. The NEP(28N–158N, 1258–1008W) and SEP (158–28S, 1258–1008W) are designated by solid–red and dashed–blue curves, respectively. The vertical axis spans 200 W m22 for all panels.

has been subtracted from the net heat flux (Fig. 11a) toremove the offset. For the net heat flux, the SEP has
a distinct annual cycle that is larger in amplitude andmore sinusoidal in shape than in the NEP. Similar north–south contrast is also found in the absorbed shortwave
flux (Fig. 11b). The absorbed shortwave flux stays smallerin the NEP than in the SEP, as found in section 3b. While
the SEP is exposed to a well-defined annual cycle ofshortwave flux, as readily expected from the seasonal
migration of insolation, seasonality is rather muted in the
NEP. Reasons for this contrast will be discussed shortly.
Although the sinusoidal seasonality of the SEP shortwave
flux resembles the net heat flux shown in Fig. 11a, theabsorbed shortwave flux only accounts for approximately
half of the amplitude. The other half can be attributed toother heat fluxes, among which the advective heat flux
appears to have the largest amplitude of seasonal cycle(Fig. 11c). The advective heat flux exhibits a clear annual
cycle with an amplitude of ;50 W m22 in the SEP anda slight phase delay relative to the shortwave flux. The
advective heat flux in the SEP reaches a maximum fromMarch through April when the SEP warm band develops(see section 4). The NEP shows a weaker but discernible
annual cycle in advective heat flux, probably due to theseasonal meridional shift of the east Pacific warm pool

(Wang and Enfield 2001). The remaining heat fluxes, allshown in Fig. 11d, do not have a notable seasonal cycle,
except that QLH in the SEP slightly varies seasonally,reaching itslargest negative flux in austral winter.
The contrast in absorbed shortwave flux between the
SEP and NEP stems from a combination of multiple effects. As illustrated in Fig. 12a, the clear-sky solar flux at
the BOA, F Yclr
SW , is almost equivalent on either side of theequator when shifted in time by a half year; however, thesmall but finite eccentricity of the terrestrial orbit leads toa smaller amplitude annual cycle in the NEP relative to
the SEP. The earth passes through perihelion in boreal
winter in the present-day orbit, reducing the northernseasonal cycle of insolation and amplifying the southern
cycle to a small extent. This tiny difference in amplitudeof the annual cycle is enhanced when the mixed layer absorptivity, increasing southward across the equator in thetropical east Pacific (Fig. 4f), is also considered. The resulting annual cycle of absorbed shortwave flux, evaluatedin the absence of any cloud-shielding effects, has an amplitude nearly twice as large in the SEP than in the NEP.
The seasonal cycle is further modulated when clouds
are included in the shortwave budget (Fig. 12b). In the
NEP, the cloud shortwave impact, or the difference between the all-sky and (hypothetical) clear-sky shortwave

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22
FIG. 12. Climatological annual time series of (a) clear-sky BOA shortwave flux (F Yclr
SW ) (W m ) and absorbed
22
22
clr
clr
shortwave flux (QSW ) (W m ), (b) shortwave cloud impact (QSW QSW ) (W m ), (c) high cloud fraction (lines,
labeled on left axis) and deep storm coverage (%) (shaded histogram, labeled on right), and (d) low cloud fraction.
The NEP (28N–158N, 1258–1008W) and SEP (158–28S, 1258–1008W) are designated by solid–red and dashed–blue
curves, respectively. The histogram is shaded in yellow (light blue) for the NEP (SEP).

fluxes, seasonally fluctuates between 240 (boreal winter)and 270 W m22 (summer). This variability is mainly
controlled by the high cloud fraction, which stays large
year round but moderately undulates loosely in phasewith deep convective activity in the NEP (Fig. 12c). The
high cloud variability there is nearly perfectly out ofphase with the solar cycle, leading to a damping effect on
the seasonality of absorbed shortwave flux in the NEP.
Such a damping effect, on the other hand, is not expected
for the SEP. Figure 12c shows that, since the period of
active convection in the SEP is limited to austral fall, thehigh cloud fraction there drops by 50% throughout theremainder of the year. On the contrary, the low cloud
fraction slightly increases from austral fall through winterin the SEP while remaining almost constant in the north.
In summary, the north–south contrast in the annual cycleof clear-sky absorbed shortwave flux is somewhat enhanced when the cloud radiative impact is included, as
the seasonal oscillation of cloudiness damps the shortwave variability in the NEP but not in the SEP.

b. Synthetic SST
The north–south difference in the annual cycle ofthe ocean surface heat flux is crucial for understanding

persistence of the north ITCZ. This is demonstrated inthis section by the synthetic SST generated by the heat
flux forcing estimated in this work. The overall strategy
follows that outlined by Masunaga and L’Ecuyer (2010)
with a few updates.
Three experiments were set up to compute the SST
with the heat budget in Eq. (1), where all heat fluxes are
given as the climatological annual time series. Besidethe control experiment, a second experiment is carriedout where the NEP flux, shifted in time by a half year, ismirrored onto the SEP; that is, Qnet(x, y, t) is replaced
by Qnet(x, 2y, t 2 T/2) for negative latitudes (y , 0), while
Qnet remains unchanged in the NEP. Similarly, a third
experiment is performed with the NEP flux replaced by
the SEP flux with the phase adjusted. The second and third
experiments, named the mirrored SEP and mirrored NEP
experiments, respectively, are aimed at finding out how thenorth and south ITCZs are altered if the seasonal cycle ofheat flux forcing is symmetric about the equator. All threeof these experiments are initialized with the observed climatological SST on 1 January and integrated over timeusing a daily time step through the end of the year. The
hQresi term on the rhs of (1) assures that the climatological
SST should be annually periodic.

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FIG. 13. (a) Climatological annual time series of observed SST (C8) in the time–latitude domain averaged over
1258–1008W and (b) annual climatology of observed SST (C8). (c),(d) As in (a),(b) but for synthetic SST from thecontrol experiment. (e),(f) As in (a),(b) but for synthetic SST from the mirrored SEP experiment. (g),(h) As in (a),(b)but for synthetic SST from the mirrored NEP experiment. See text for experimental design. The study area in (a),(c),(e),and (g) is indicated by the dashed box.

The experimental results are shown in Fig. 13 together
with the observed SST for reference. The time–latitudesection of observed SST (Fig. 13a) clearly delineates thepersistent north ITCZ with a subtle seasonal north–south migration and the SEP warm band temporarily
developing in austral fall. In the control experiment,
shown in Fig. 13c, the north ITCZ is split into two bands
from April to July. This artifact, arising in the vicinity ofthe equator, is probably due to the limited capability of
estimating the upwelling heat flux, as noted in section
3a. Outside the equatorial band, the control experiment
overall reproduces the observed properties. When averaged

over the year, the control experiment is reasonably
close to the observed geographical SST distribution(Figs. 13b and 13d).
No sign for the austral fall double ITCZ is identifiablein the mirrored SEP experiment (Figs. 13e and 13f). Notethat a 1808 phase difference in the heat flux seasonality
between the SEP and NEP resulted in an artificial discontinuity at the equator. In this experiment, the seasonal
cycle of the net heat flux is weakened over the SEP and is
no longer able to give rise to the SEP warm band. The
meridional SST contrast introduced by the initial conditions is more or less preserved throughout the year in the15 MARCH 2011

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1797

absence of a strong seasonal variation in heat flux forcing.
In contrast, the mirrored NEP experiment provides a very
different result. Synthetic SST from this experiment is
now seasonally variable in the NEP as well as in the SEP,
leading to a seasonally alternating ITCZ appearing oneither side of the equator (Fig. 13g). The enhanced seasonality in the NEP no longer favors a persistent ITCZ.
The SST in the NEP warms to 298C in boreal fall but thiswarming is nearly canceled out when averaged over theyear. The north ITCZ, as far as inferred from the synthetic SST, almost vanishes between January and Augustin the mirrored NEP experiment (Fig. 13h). Thus, it canbe inferred that the persistence of the north ITCZ islargely caused by the damped seasonal cycle of the NEPheat flux forcing.

6. Discussion and conclusions
This paper addresses a long-standing problem with theequatorial asymmetry of the east Pacific ITCZ from theperspective of the ocean surface heat budget. While coupled model simulations have been successfully exploitedby many previous studies relevant to this problem, thecurrent analysis makes use of satellite measurements forevaluating the heat flux forcing to the ocean mixed layer.
The annual mean climatology of the absorbed shortwave
flux shows a prominent meridional asymmetry, marked
with a 40 W m22 difference between 108S and 108N, mainly
due to shortwave reflection by high clouds associatedwith ITCZ convection. The mystery does not lie in whatsustains the equatorial asymmetry against the annuallysymmetric solar forcing as is often claimed. A proper way
to pose the problem is how the asymmetric shortwave
forcing is counteracted by other heat fluxes in tandem andhow the resulting heat flux distribution physically regulates the ITCZ structure, including a possible feedback tohigh cloud variability responding to ITCZ convection.
Xie and Seki (1997) analyzed the individual contributions of wind, humidity difference, cloudiness, and sea–airtemperature difference to the ocean surface heat budget
and concluded that the sun-shielding effect of high clouds
is largely offset by high relative humidity in the equatorial
asymmetry. They argued that the meridional asymmetryof wind is crucial for maintaining the observed north–south asymmetry of SST in the way that latent heat flux
as a whole is locally balanced with the incoming radiativeflux. An analysis similar to that of Xie and Seki (1997) is
performed to see if the same conclusions are drawn fromour surface flux estimates (Fig. 14). Figure 14 shows theequatorial asymmetry in latent heat flux and absorbed
shortwave flux as in Fig. 3 and two additional hypotheticallatent heat fluxes computed with each of the wind speed
and humidity difference individually made symmetric about

FIG. 14. North–south difference in climatological zonal-mean(1258–1008W) latent heat and absorbed shortwave fluxes, computed for each latitude denoted by the abscissa. North exceeds
south when the differential flux is positive. Plotted are latent heat
flux (QLH, solid) as it is; QLH recomputed with an adjusted humidity difference; ra[q*y (Ts) 2 qy], symmetric about the equator(dashed); and QLH recomputed with an adjusted wind speed
symmetric about the equator (dotted–dashed). The shaded histogram shows the absorbed shortwave flux as in Fig. 3.

the equator in (2). The equatorial asymmetry in latent
heat flux would be somewhat enhanced (reversed) if
the humidity difference (wind speed) were symmetric,
as one might expect from Fig. 5. The expected magnitude of the asymmetry in latent heat flux by perturbing
the humidity difference and wind speed, however, ismuch too small to fully offset the north–south contrast
in absorbed shortwave flux at ITCZ latitudes, unlike
the findings of Xie and Seki (1997). The origins of thequantitative disagreement from this earlier work are
unknown, but are possibly due to the fact that the Xieand Seki analysis relies on an empirical parameterization to evaluate shortwave flux rather than a direct heat
flux estimate.
The asymmetric shortwave forcing is, in fact, balanced
not by the latent heat flux alone but by the combined effects of several heat fluxes. Among those participating in
the east Pacific heat budget, the advective heat flux makesa larger contribution to the equatorial asymmetry than allother fluxes except the shortwave flux. This may be a little
surprising because the advective heat flux is least important in magnitude for the east Pacific heat budget as awhole. The annual mean advective heat flux changes sign
to positive as latitude increases beyond 78N, owing to thewarm water transport from the east Pacific warm pool and,to a lesser extent, the NECC. The meridional asymmetryarises because the SEP has neither a warm pool nor an
equatorial countercurrent to supply warm water. The latent
heat flux, largest among the negative (or cooling) fluxes,

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has a moderate asymmetry governed by the scalar windspeed that is somewhat larger in the SEP than in the NEP.
The longwave flux is relatively homogeneous across theeast Pacific. The upwelling heat flux is largest in bothmagnitude and asymmetry within a narrow equatorial
band between 58S and 58N while it quickly diminishes
poleward beyond 58S/N. The wind–upwelling interactionthat amplifies the equatorial asymmetry, as suggested by
Mitchell and Wallace (1992) and Chang and Philander(1994), may be operating in this limited latitudinal range
centered on the equator, but is unlikely to directly influence the whole ITCZ structure, as also pointed out by
Xie (2004).
A key to understanding the ITCZ asymmetry lies in thenorth–south contrast of the seasonal cycle in combination
with the annual mean field. The seasonal cycle of heat flux
forcing is found to be muted in the NEP relative to the
SEP. Three separate factors account for this difference: 1)
annual fluctuation of the earth–sun distance, 2) meridional
gradient in mixed layer absorptivity, and 3) seasonal variation of the cloud shortwave impact. The first factor was
also pointed out by Wang and Wang (1999) but in a somewhat different context. The third point does not have aslarge an impact as the first two combined but has noteworthy implications. High cloudiness tends to increase
as deep convection intensifies from a cold to a warmer
season, masking, in part, the seasonality of the BOA
shortwave flux in the NEP. This may be viewed as aseasonal analog of the thermostat feedback (Ramanathan
and Collins 1991), although a close examination of theunderlying mechanism is outside the scope of the present
study. On the other hand, maritime stratocumuli, dominating the low cloud fraction over the SEP, prefer colder
oceans and potentially reinforce the preexisting seasonaland spatial variations of SST, as suggested by Philander
et al. (1996). Such a low cloud impact on the shortwave
flux is found to exist in the SEP but not to an extent thatwould overwhelm the high cloud impact in either theannual mean or seasonal cycle. The main body of thesoutheast Pacific stratocumulus cloud deck is indeed eastof the current study area, closer to the Peruvian coast (seeFig. 4d). The stratocumulus regulation of SST, if at work,
may be a localized effect near the coast, but a cold SST
there could affect a broader region far to the west throughadvection. The advective heat flux, a major ingredientof the equatorial asymmetry, originates from a remarkable contrast between the east Pacific warm poolin the north and a cloud-covered cold pool in the southproviding a pathway for continental geography to impact the ITCZ asymmetry (Philander et al. 1996; Xieand Saito 2001). Regions under the direct influence ofthe north–south SST disparity caused by the combination of these geographical and stratocumulus effects

VOLUME 24

can be communicated by advective flux to the rest of theeastern Pacific basin.
The month-to-month variation of wind speed and latent
heat flux in the SEP is found to be positively correlatedwith north ITCZ activity as expected from the WES
feedback hypothesis. The WES feedback, on the other
hand, is known to lose its control on the ITCZ asymmetrywhen exposed to a strong seasonal forcing (Xie 1996,
1997a). Since the asymmetric WES mode grows at a rate
much slower than a seasonal time scale (Xie 1997b), theability of the WES feedback to explain the observed ITCZ
asymmetry is limited by the overwhelming seasonality ofsolar forcing. Xie (1996, 1997a) demonstrated that the east
Pacific ITCZ would become more equatorially symmetricin the annual mean as the seasonal forcing strengthens,assuming that the amplitude of seasonal shortwave forcing is equal on both sides of the equator. In reality, however, the seasonal cycle of the absorbed shortwave flux has
substantially different amplitudes between the Northernand Southern Hemispheres. As a result, the ITCZ retains
its equatorial asymmetry because the magnitude of seasonal forcing is a critical element for the persistence ortransience of the ITCZ.
A set of simple experiments were carried out to study
the role of the seasonally varying heat flux forcing in thepersistence of the north ITCZ. Modeled SST is found tobe subject to a strong seasonal cycle in the north ITCZ
if the NEP heat flux forcing is replaced by SEP values,
leading to a significant weakening of the north ITCZ in
the annual mean. Interestingly, the resultant seasonality
when the north and south ITCZs emerge and disappear inturn is similar to the ‘‘seasonally alternating ITCZ error’’inherent in some coupled climate models (de Szoeke andXie 2008). The regional contrast in the seasonality of heat
flux forcing might be one of the factors responsible forsuch model biases.
This paper is focused on the ocean mixed layer headbudget and its potential impact on the atmospheric dynamics has yet to be examined further in depth. The
current work needs to be extended to future studies withthe complete circle of air–sea feedback considered to fully
understand the interactive mechanisms sustaining the
ITCZ. In addition, the heat budget analysis conductedhere relies on many assumptions and simplifications thatcompensate relatively limited pieces of information available from satellite observations. The present results andimplications would need verification by coupled ocean–
atmosphere models of reasonable complexity and flexibility where the processes vital for the ITCZ asymmetry
are individually tested. Such modeling studies exist but
have more or less focused on assessing the effects of wind,
upwelling, and low clouds with known hypotheses in mind,
where some key processes such as advective heat flux have

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been largely overlooked. Benchmark estimates of the heat
fluxes and resulting cloud structure as provided by thisstudy could be used to initialize and/or evaluate futureclimate model simulations, which would lead to a deeper
understanding of the equatorial asymmetry of the east
Pacific ITCZ.
Acknowledgments. The TRMM PR (2A25) and VIRS
(1B01) datasets were provided by the Japan Aerospace
Exploration Agency (JAXA) Earth Observation ResearchCenter (EORC), and the TMI- and AMSR-E-retrieved
SST and CWV datasets by the Remote Sensing Systems(information online at www.ssmi.com). The QuikSCAT
SeaWinds ocean wind vector data and OSCAR surface
current product were obtained from the Physical Oceanography Distributed Active Archive Center (PO.DAAC)
at the NASA Jet Propulsion Laboratory, Pasadena, California (information online at podaac.jpl.nasa.gov). The
AIRS level 3 daily gridded product was acquired as part of
the activities of NASA’s Science Mission Directorate, andis archived and distributed by the Goddard Earth Sciences(GES) Data and Information Services Center (DISC).
The WOA05 temperature data were obtained from the
NOAA/National Oceanographic Data Center. The mixed
layer depth dataset was downloaded online (www.lodyc.
jussieu.fr/;cdblod/mld.html). This work is supported by
the Ministry of Education, Culture, Sports, Science, andTechnology (MEXT) Grant-in-Aid for Young Scientists
20740269. The HERB dataset was generated under the
support of National Aeronautics and Space Administration (NASA) Energy and Water Cycle Study (NEWS)
Grant NAG06GC46G.
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