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Summer Sea Surface Temperature Conditions in the North Atlantic and Their Impact
upon the Atmospheric Circulation in Early WinterCHRISTOPHE CASSOU
Climate Global Dynamics Division, NCAR, Boulder, Colorado, and Climate Modelling and Global Change Team, SUC, CERFACS,
Toulouse, France

CLARA DESER
Climate Global Dynamics Division, NCAR, Boulder, Colorado

LAURENT TERRAYClimate Modelling and Global Change Team, SUC, CERFACS, Toulouse, France

JAMES W. HURRELL
Climate Global Dynamics Division, NCAR, Boulder, Colorado

MARIE DRÉVILLON
Climate Modelling and Global Change Team, SUC, CERFACS, Toulouse, France(Manuscript received 26 February 2003, in final form 3 February 2004)
ABSTRACTThe origin of the so-called summer North Atlantic ‘‘Horseshoe’’ (HS) sea surface temperature (SST) modeof variability, which is statistically linked to the next winter’s North Atlantic Oscillation (NAO), is investigatedfrom data and experiments with the CCM3 atmospheric general circulation model (AGCM). Lagged observational
analyses reveal a linkage between HS and anomalous rainfall in the vicinity of the Atlantic intertropical convergence zone. Prescribing the observed anomalous convection in the model generates forced atmospheric Rossby
waves that propagate into the North Atlantic sector. The accompanying perturbations in the surface turbulent
and radiative fluxes are consistent with forcing the SST anomalies associated with HS. It is suggested that HS
can therefore be interpreted as the remote footprint of tropical atmospheric changes.
The ARPEGE AGCM is then used to test if the persistence of HS SST anomalies from summer to late fall
can feed back to the atmosphere and have an impact on the next winter’s North Atlantic variability. Observed
HS SST patterns are imposed in the model from August to November. They generate a weak but coherent early
winter response projecting onto the NAO and therefore reproduce the observed HS–NAO relationship obtained
from lagged statistics. Changes in the simulated upper-level jet are associated with the anomalous HS meridional
SST gradient and interact with synoptic eddy activity from October onward. The strength and position of thetransients as a function of seasons are hypothesized to be of central importance to explain the nature, timing,and sign of the model response.
In summary, the present study emphasizes the importance of summer oceanic and atmospheric conditions inboth the Tropics and extratropics, and their persistence into early winter for explaining part of the NAO’s lowfrequency variability.

1. Introduction
Over the last couple of years, interest in the lowfrequency variability of North Atlantic climate has increased markedly (e.g., Marshall et al. 2001; Stephenson
Corresponding author address: Dr. Christophe Cassou, CERFACSSUC (URA 1875), Climate Modelling and Global Change Team, 42
Ave. Gustave Coriolis, 31057 Toulouse, France.
E-mail: cassou@cerfacs.fr

q 2004 American Meteorological Society

et al. 2003). In particular, attention has been focused on
the North Atlantic Oscillation (NAO), described as themost prominent and recurrent pattern of atmosphericvariability (Wallace and Gutzler 1981). Important ecological, economical, and societal impacts that the NAO
exerts on the entire Northern Hemisphere have been
documented and questions have been raised about its
potential predictability (Hurrell 2003). Although the
NAO mainly arises from internal atmospheric interac-

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tions, external forcings such as sea surface temperature(SST) anomalies or increasing greenhouse gases, etc.
might affect the occurrence and/or the persistence of its
phase. Their influence could be important for understanding both amplitude and long-term evolution of the
NAO in the context of climate change or climate predictability.
On one hand, the impact of atmospheric circulation
anomalies upon extratropical/subtropical oceans has
been demonstrated in numerous studies (see, e.g., Battisti et al. 1995). Anomalous turbulent heat fluxes, buoyancy-driven entrainment and Ekman advection are themain processes whereby the atmosphere imprints largescale SST anomalies. In such a scenario, the extratropical atmosphere primarily leads the midlatitude ocean
variability by a month or so, as examined for instancein Deser and Timlin (1997) or Frankignoul et al. (1998).
The low frequency oceanic anomalies are considered as
the passive temporal integration of the atmospheric stochastic forcing (Frankignoul and Hasselmann 1977).
On the other hand, although the atmospheric fluctuations are undoubtly the dominant driver of the upperocean anomalies, the role of the ocean has been suggested to explain the reddening of the atmospheric fluctuations from seasonal to decadal time scales. For instance, Rodwell et al. (1999) and Mehta et al. (2000)
using atmospheric global circulation models (AGCMs)
forced by historical global SST estimations have beenable to reproduce some of the low-frequency variabilityof the winter NAO during the last 50 years. The dominant influence of the tropical/subtropical Atlantic oceanic basin has been inferred in some studies (e.g., Sutton
et al. 2001). Terray and Cassou (2002) diagnose frommodel simulations that anomalous SSTs there alter the
local atmospheric Hadley cell over the western tropicalbasin, providing an anomalous source of Rossby waves(Sardeshmukh and Hoskins 1988) extending northeastward from the Caribbean toward Europe. The phase of
the NAO is ultimately affected by this teleconnection
mechanism as described in Peng et al. (2003), Drévillon
et al. (2003), and Cassou et al. (2004). Hoerling et al.
(2001) argued that the whole Tropics, not just the Atlantic domain, must also be considered. In particular,
the role of the Indian Ocean has been suggested by
Sutton and Hodson (2003) and Hurrell et al. (2004).
North Atlantic midlatitude SST anomalies may also
have a nonnegligible influence on the variability of the
NAO. The dominant mechanisms are discussed in Pengand Whitaker (1999) and Peng and Robinson (2001).
They emphasize the role of synoptic storm activity andassociated eddy–mean flow interaction that can be perturbed by the existence of anomalous meridional SST
gradients along the subpolar–subtropical gyre front. Extratropical oceanic anomalies may also contribute to thereddening of the NAO spectrum in the interannual bandthrough the reemergence mechanism in which upperocean temperature anomalies created over a deep mixed
layer in winter are preserved within the highly stratified

VOLUME 17

summer thermocline (Deser et al. 2003). As they reappear at the surface at the next late fall or early winterby reentrainment (de Coëtlogon and Frankignoul 2003),
they may have a feedback impact upon the overlying
atmosphere. However, conclusions about extratropical
SST forcing upon the North Atlantic atmosphere arevery model dependent [see Robinson (2000) or Kushnir
et al. (2002) for a review] and should be taken withcaution. The models’ different sensitivities appear to be
due to a different representation of the simulated climatological flow and internal atmospheric modes of variability.
At monthly to seasonal time scales, lagged relationships between the early wintertime North Atlantic atmosphere and the previous summer–fall oceanic conditions have been documented by Czaja and Frankignoul(1999, 2002) and Drévillon et al. (2001). Using singularvalue decomposition (SVD), a statistical link is extracted between the so-called summer North Atlantic‘‘Horseshoe’’ (HS) SST pattern and the NAO, as illustrated in Fig. 1 using data from the National Centers
for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis (Kalnay
et al. 1996). The negative phase of the NAO (Fig. 1a)
in early winter [October–December (OND) mean] ischaracterized by a slackening of the Icelandic low and
the Azores high, and is associated with late summer[August–September (AS) mean] negative SST anomalies located southeast of Newfoundland and extending
westward slightly along the Gulf Stream path, surrounded by positive anomalies on the eastern side of the Atlantic (Fig. 1b). Maximum warming occurs off thenorthern African coast and from Britain to the southern
tip of Greenland. The HS mode is dominated by interannual fluctuations with a tendency toward dominantpositive phases in the 1950s and 1960s and more recurrent negative phase from the early 1970s to the early
1990s (Fig. 1c). Watanabe and Kimoto (2000) and Drévillon et al. (2001, 2003), from data and model experiments respectively, speculate on physical processes responsible for the statistical lagged properties shown in
Fig. 1, but the nature of the summer ocean influence on
the next winter NAO is still unclear.
The main goal of this paper is to go beyond statisticalanalyses and to suggest a physical explanation for the
existence of the summer HS pattern and a hypothesis
for the lagged relationship between that SST pattern andthe following winter atmospheric conditions. The observational results are used to motivate several modelexperiments to test these ideas. The paper is organized
as follows. A combined observation–model analysis ispresented in section 2 to illustrate the relationship between anomalous summer atmospheric conditions over
the Atlantic and the oceanic HS mode. Interpretations
about the origin of the HS are discussed. The lagged
influence exerted by the HS on the North Atlantic wintertime atmosphere is then examined in section 3. The
observed HS monthly evolution is prescribed from late

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FIG. 1. First singular value decomposition modes calculated between (a) OND SLP (hPa) and (b) AS SST (8C) computed from
NCEP–NCAR over the 1949–2001 period. The covariance between
the two modes is equal to 63.9%. Contour intervals are respectively
0.1 hPa and 0.18C. (c) Normalized time series related to the AS SST(bars) and OND SLP (solid line) leading SVD mode. The correlation
between the two PC time series reaches 0.72 (significant at 99%).

summer to early winter in a model to investigate its
impacts on the North Atlantic atmospheric dynamics. A
synthesis and discussion are provided in section 4.
2. Origin of the late summer North Atlantic‘‘Horseshoe’’ SST modea. Motivation: Observational results
The origin of the late summer HS mode that exhibits
the strongest covariability with the next winter NAO

3351(Fig. 1) is investigated first from observations. To clearly identify the role of the atmosphere in the origin of
the ocean anomalous pattern, June–July (JJ) averaged
atmospheric fields have been regressed onto the AS averaged SST SVD principal component (PC) time series.
All atmospheric fields have been quadratically detrended prior to the regression.
There is some evidence that the extratropical HS
mode is forced in part by atmospheric circulation anomalies through modifications of the surface turbulent andradiative fluxes. The lagged regression of the total surface heat flux estimated from the Comprehensive
Ocean–Atmosphere Data Set (COADS; Woodruff et al.
1998) following the Cayan (1992) formulation matches
the spatial pattern of the HS SST anomalies fairly well,with the atmosphere tending to warm the subtropicaland eastern Atlantic and to cool the western basin (Fig.
2a).
The turbulent flux contribution can be mainly understood in terms of near-surface wind changes (Fig.
2b). The anomalous cyclonic circulation located off
Newfoundland induces anomalous northerlies over thewestern part of the basin, slackened westerlies at higher
latitudes especially in the Labrador and Irminger Seas,and reduced trade winds within the subtropics in thecentral Atlantic. Cold air is advected from the north on
the western side of the basin, cooling the ocean mixed
layer, while decreased winds within the band 508–658N
reduce the turbulent heat flux, warming the mixed layer.
Besides, Ekman pumping due to the anomalous cyclonic
circulation over the basin may contribute to cooling
down the ocean surface off Newfoundland (not shown).
Advection is also expected to explain part of the anomalies along the Gulf Stream path where local wind anomalies are not clearly related to surface oceanic conditions.
The trade winds are not significantly modified in theeastern subtropical basin where the maximum warming
occurs. Over this specific region, the radiative balanceat the surface might have a complementary or evengreater role. Decreased cloudiness is observed on theeastern side of the Atlantic basin especially within thesubtropical band and off Europe (Fig. 2c), consistent
with the higher SSTs. Reduced low-level clouds that are
dominant over this area diminish the so-called albedo
feedback and tend to warm up the surface ocean. Increased cloudiness over most of the anomalous SST coldcore reinforces the cooling of the ocean mixed layer due
to enhanced turbulent heat loss. Although less significant and less large scale, the NCEP–NCAR surface
shortwave flux anomalies match the COADS cloudinessresults. The cloud pattern is persistent until fall (notshown) and may contribute to maintain the ocean surface anomalies. This is particularly true above the subtropical SST lobe. The western tropical Atlantic basin(west of 608W) is characterized by increased cloudinessover positive SST anomalies. Low-level stratiform
clouds are fairly sparse over this area dominated by

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convection, and this pattern suggests some strengthening of the convective activity.
We investigate this signal further using NCEP–NCAR
outgoing longwave radiation (OLR) as a proxy for deep
convection in the Tropics. As accurate measurements
have been only available from satellites since the late1970s, station-based precipitation data (Hulme and New
1997) are used to further confirm the robustness of ourfinding from OLR. The HS mode is associated with earlysummer enhanced convective activity (decreased OLR)
over Africa and over the Caribbean Sea extending northwestward toward Florida (Fig. 2d). Such an increase of
convection within the band 108–208N, indicative of aslight northward shift and intensification of the intertropical convergence zone (ITCZ), contrasts with mostly
dry conditions over the Amazonian basin (increased
OLR). The precipitation station data confirm the OLR
signal over land (Fig. 2d) and are consistent with increased cloudiness over the Caribbean diagnosed independently from COADS. The anomalous rainfall pattern shown in Fig. 2d is similar to that documented by
Giannini et al. (2003) at interdecadal time scale and by
Sutton et al. (2000) and Okumura et al. (2001) at interannual time scales. As shown by Fontaine et al.
(1999), when the SST anomalies are positive in thenorthern tropical Atlantic, the local Hadley circulationis altered with increased low-level convergence from
Gulf of Guinea to the Sahel bringing more moisture
over land consistent with enhanced convection. By contrast, mostly dry conditions dominate along the equator
over sea and over the South American continent.
b. CCM3 model experiments

FIG. 2. Two-month lead (JJ) regression onto the AS SST PC of (a)total surface heat flux (W m 22 ) from COADS and (b) 1000-hPa
NCEP–NCAR wind superimposed on the SVD SST mode itself. Contour intervals are 0.18C. (c) Cloudiness (octas) from COADS and (d)OLR (W m 22 ) (contour) from NCEP–NCAR and rainfall stationbased data (mm day 21 ) from Hulme and New (1997) (contour interval
is 1 W m 22 for OLR and positive values are indicative of decrease
convection). Only points satisfying the 90% level confidence are plotted for all the datasets.

We have shown from observations that the late summer North Atlantic SST anomalies associated with the
HS mode are likely a lagged local response to the extratropical underlying atmospheric changes. We havealso shown that HS is preceeded by anomalous tropicalatmospheric anomalies. To explore the extent to which
tropical Atlantic rainfall changes can affect the midlatitude atmosphere in summer and how these are linked
to HS, two experiments are conducted with the third
version of the NCAR Community Climate Model(CCM3), which has a T42 triangular horizontal resolution and 18 vertical levels. A complete description of
the physical and numerical methods used in CCM3 isprovided in Kiehl et al. (1996) and model validation is
given in Hurrell et al. (1998). The summer ITCZ intensity is artificially perturbed in the model based upon
the observed anomalous convection pattern associatedwith HS (Fig. 2d). Two idealized heat sources have been
specified, one centered over the Caribbean Sea and one
over the Sahel region (rectangles in Fig. 3a). The vertical distribution of the heating perturbation (Fig. 3b)reaches a maximum intensity at 500 hPa and is equal
to 128C day 21 , which roughly corresponds to twice
the maximum observed changes estimated from the1 SEPTEMBER 2004

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F IG . 3. (a) The JJ 2-month-averaged difference for OLR(W m 22 ) between anomalous prescribed diabatic ‘‘heating’’ and diabatic ‘‘cooling’’ CCM3 experiments. Contour interval is 10 W m 22
and shading stands for positive anomalies. Convention is negative
OLR for increased convection. The thick solid boxes represent thegeographical domains where the anomalous diabatic heating/coolingis prescribed in CCM3. (b) Anomalous diabatic heating vertical profile (8C day 21 ) for a section at 158N as a function of longitude andprescribed in CCM3. Contour interval is 0.58C day 21 .

NCEP–NCAR reanalyses. A twin experiment with aprescribed 228C day 21 is also carried out and the modelis integrated for 10 summers in both cases. The reader
is invited to refer to Branstator and Haupt (1998) for acomplete description of the experimental setup. The
choice for the strong magnitude of the perturbation applied in the model is dictated by our attempt to better
extract the tropical impact onto the midlatitude dynamics within the constraint of limited computing resources.
In the following, we focus our explanation on the linear[(heating 2 cooling)/2] response.
As expected, the model shows enhanced convective
activity collocated with the prescribed anomalous diabatic heating (Fig. 3a). Maximum OLR anomalies reachabout 260 W m 22 centered over the Sahel and Caribbean. Diminished convection appears over the SouthAmerican continent and the western equatorial Atlantic,
reminiscent of the observations. Thus, the diminishedconvection may be interpreted as a remote response to
the increased convection along the ITCZ in the model.
We have also verified that specifying a negative heatsource over Amazonia in the model leads to enhanced
convection over the Caribbean and also to a lesser extent

FIG. 4. As in Fig. 3a, but for CCM3 (a) wind at 1000 hPa, (b)geopotential height at 500 hPa (m) and (c) total heat flux (W m 22 ).
Contour intervals are 5 m and 10 W m 22 . Shaded areas exceed the
95% significance limit using T statistics.

over the Sahel (not shown). This suggests that the fullheating distribution in the whole tropical Atlantic may,in fact, play a role. Very similar conclusions are obtained from the upper-tropospheric convergence field(not shown).
The simulated low-level wind response to the heatsources is shown in Fig. 4a. Weakened trade winds dominate the subtropical North Atlantic basin, consistent
with observations. Similarities can be also found between the modeled and observed surface wind changes
at midlatitudes. Despite a slight westward and northward

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shift, the model is able to reproduce the cyclonic atmospheric circulation off Newfoundland in association
with a stronger Atlantic ITCZ. The perturbed atmospheric circulation is reminiscent of a Rossby wave train
pattern originating from the Gulf of Mexico and extending northeastward toward Scandinavia. Althoughsuch a mechanism is more active in winter–spring (Webster 1982), it seems to be at work also in summer inthe model. It clearly dominates the remote midlatitude
atmospheric changes as displayed in Fig. 4b for the 500hPa geopotential height (hereafter Z500). Minimum
Z500 is located around Newfoundland and slightly westof the low-level cyclonic conditions while high pressure
over Britain is consistent with the surface anticyclone.
Note that the latter is less evident in the observations(Fig. 2b). The overestimation of the Rossby wave extension simulated in CCM3 might be attributed to thehigher intensity of the perturbation prescribed in the
Tropics compared to observations or to the model oversensitivity for that particular season due to mean state
biases. To measure the importance of the remote signal,
we note that the signal-to-noise ratio comparing forced
variability (model response in sensitivity experiments)
and internal variability (estimated from a control simulation of CCM3) is close to 1 off Newfoundland anddrops to 0.6 over the United Kingdom.
To further confirm the teleconnection hypothesis andits origin, complementary simulations have been carriedout by selecting only one of the two anomalous diabatic
heating sources, thereby isolating the Sahel versus Caribbean impacts. Similarly to the previous experiments,
the model is integrated for 10 summers and for the‘‘heating’’ and the ‘‘cooling’’ cases. The Rossby-wave
type anomalous circulation completely disappears in the
Sahel-only experiment while a clear cyclonic circulation
centered off the North African coast (308N, 208W) dominates (not shown). In that case, the slackening of theAtlantic trade winds is more pronounced since the lowlevel convergence to the Caribbean has been removed.
By contrast, Caribbean-only experiments (not shown)
reveal a clear Rossby wave train while the trade winds
are reinforced in the western tropical basin but not affected in the eastern basin. This result is consistent withthe fact that the western convection core is bordered on
its northern side by climatological upper-level westerlies, whereas the African core is capped by easterlies,
precluding any meridional propagation of tropicalforced Rossby waves (Hoskins and Pierce 1983). Basedon these model results, we suggest that the observed
Atlantic wind pattern associated with the HS in summer
can be viewed as a remote response to changes in tropical convection over both the Sahel and Caribbean regions, with the trade winds responding to both heatsources and the extratropical winds responding mainly
to the Caribbean one.
Total surface heat flux anomalies over the subtropical/
North Atlantic basin in the CCM3 experiments represent
the local atmospheric forcing of the ocean mixed layer

VOLUME 17(Fig. 4c). The heat flux anomalies are consistent withanomalous low-level wind conditions and bear a strong
resemblance to the observed pattern (Fig. 2a). Weaker
trade winds are associated with diminished turbulent
heat flux out of the ocean east of 508W, which tends towarm up the subtropical North Atlantic. A secondary
core centered around 508N off Europe corresponds tothe midlatitude slackening of the mean westerlies andsimilarly tends to warm up the northeast Atlantic. By
contrast, the western side of the basin north of 208N isdominated by enhanced evaporation and northerly
winds that tend to cool down the surface ocean, althoughsignals are rather weak and mostly confined south of
Newfoundland close to the Gulf Stream. Decomposing
the total heat flux into radiative and turbulent components (not shown) reveals that the latter dominates in
the model except off the Iberian penisula where thesimulated warming tendency appears to be limited tothe coast (Fig. 4c) compared to observations (Fig. 2a).
This confinement can be attributed to an overestimationof enhanced cloudiness in the model over this area (theCCM3 cloud response to the perturbed ITCZ is maximum to the east of the anomalous cyclonic circulation
while in reality these are collocated, Fig. 2c). Except
over this specific area, changes in radiation balance due
to cloud cover are weak in the model and are suspected
to be underestimated. This could be explained by theabsence of local ocean feedback to the atmosphere dueto the experimental configuration. In reality, local SST
anomalies have been shown within the subtropics tolocally affect the stability of the atmospheric column
and thus amplify and/or maintain the anomalous cloud
pattern (Norris et al. 1998). The weakness of the cloud
impact could also be explained by the underestimationof low-level cloud amount in CCM3. Additional experiments using a newer version of CCM3 coupled to
a complex ocean mixed layer model (Alexander et al.
2002) are currently in progress to examine the SST response to tropical diabatic heating anomalies.
3. Summer SST forcing of the NAO in early
wintera. Experimental setup
The speculated feedback of SST anomalies associatedwith the HS mode onto the early winter atmosphere over
the North Atlantic–European sector is now examinedthrough two 30-member ensemble integrations, each 9months long (July–March), using the Action de Recherche Petite Echelle Grand Echelle (ARPEGE) model
jointly developed by Météo-France and European Centre of Medium-Range Weather Forecasts (Déqué et al.
1994). The present climate version of ARPEGE uses a
T63 triangular horizontal truncation. Diabatic fluxes andnonlinear terms are calculated on a gaussian grid of
about 2.88 latitude 3 2.88 longitude. The vertical isdiscretized over 31 levels (20 levels in the troposphere)1 SEPTEMBER 2004

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FIG. 5. Anomalous SST patterns (8C) added to (subtracted from) ARPEGE boundary conditions from (a) Aug to (d) Nov for the HS1(HS2) experiment. Contour interval is 0.28C and shading stands for positive anomalies.

using a progressive vertical hybrid coordinate extending
from the ground up to about 34 km (7.35 hPa). Further
details about the model physics package and about thevalidation of the simulated intraseasonal and interannual
variability can be found in Doblas-Reyes et al. (1998,
2001) and in Cassou and Terray (2001).
The HS-type SST anomalies are either added to (hereafter HS1 experiment) or subtracted from (HS2) themodel boundary conditions from 15 July to 15 December (Fig. 5), while late December to March oceanic
conditions remain unperturbed and fixed to the seasonal
climatology. The monthly anomalous SST patterns havebeen obtained by regression of monthly SST data upon
the PC time series of the leading SVD mode shown in
Fig. 1c. The intensity of the SST anomalies imposed in
the model is given by the strongest value of the PC timeseries over the analyzed 52-yr period. For HS1, maximum cooling off Newfoundland reaches 21.28C in August whereas subtropical values are slightly above
10.68C off the northwest African coast. Maximum
warming peaks at about 118C off western Europe. The
opposite is taken for HS2.
As described in Drévillon et al. (2001), the HS oceanic anomalies are fairly persistent from summer to ear-ly winter (the correlation between August and November PC time series calculated from empirical orthogonal
function decomposition of the consecutive July–August–
September–October–November normalized months andcapturing HS is equal to 0.44). Consistent with theirstudy, the extratropical lobes prescribed in our case exhibit a slow decay in amplitude with time (about 50%
diminution from August to November), while a slighteastward propagation of the midlatitude SST core isnoticeable. Compared to the extratropics, the subtropicaloceanic anomalies are more persistent both in amplitude
and space. The rate of decay of the tropical/subtropical
SST anomalies are roughly consistent with the stochastic climate model (Frankignoul and Hasselmann1977). In the following, we examine the linear [(^HS1&
2 ^HS2&)/2] component of model response (angle
brackets standing for ensemble mean).
b. ARPEGE atmospheric response to HS SST
anomalies
Zonal wind anomalies at 200 hPa (U200), Z500, andgeopotential height anomalies at 1000 hPa (Z1000) arepresented for the OND average for ARPEGE (Figs.

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FIG. 6. Regression of NCEP–NCAR OND zonal wind at (a) 200 hPa (U200) (m s 21 ), (b) geopotential height at 500 hPa (Z500) (m), and(c) at 1000 hPa (Z1000) (m), onto AS SST time series obtained from SVD. Linear ARPEGE response in OND for (d) U200, (e) Z500, and(f ) Z1000. Contour intervals are 0.5 m s 21 for U200 and 2 m for geopotential. Shaded areas exceed the 95% significance limit using
T statistics.

6d,e,f) and observations (Figs. 6a,b,c). The latter are
estimated from NCEP–NCAR and obtained by lagged
regression onto the August–September SST SVD PC
time series. The model develops a global quasi-barotropic response to the prescribed HS SST anomalies
and reproduces fairly well the observed early winteratmospheric changes. The U200 zonally elongated tripole from the Tropics to high latitudes is well captured,
especially on the eastern side of the basin. It corresponds to the reinforcement and eastward extension of
the climatological subtropical upper-level jet. The
Z500 stationary waves are accordingly modified. A
clear barotropic negative NAO signal is extracted for

NCEP–NCAR (Figs. 6b,c), and the Z1000 anomalies
are very similar in location to the MSLP structure extracted by SVD (Fig. 1a). ARPEGE geopotential height
responses (Figs. 6e,f) bear some resemblance to theobserved patterns, with generally negative (positive)
anomalies south (north) of 608N. However, the strengthand eastward extension of the positive anomalies near608–708N are significantly underestimated in the modeland the negative anomalies over the Barents Sea areopposite to observations.
A different picture emerges for the late summer (August–September) anomalies simulated by ARPEGE(Fig. 7). The Z500 NCEP–NCAR anomalies reveal a1 SEPTEMBER 2004

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FIG. 7. As in Figs. 6b,c,d,f, but for the Aug–Sep mean.

clear wavy pattern emanating from the western tropicalAtlantic basin and extending toward Scandinavia. Thispattern bears a strong resemblance to the one obtained
from the CCM3 sensitivity experiments as analyzed in
the previous section for early summer. This result suggests that the tropical forcing of the midlatitude atmosphere is persistent over the entire season and therefore
tends to reinforce the HS-type oceanic pattern initiatedat the begining of summertime. ARPEGE results contrast strongly. The model does not capture the forced
Rossby wave pattern at 500 hPa but departures between
observed and modeled signals are even more striking
in the lower troposphere. While a clear quasi-barotropic
structure is extracted from NCEP–NCAR north of 358N,
the model develops a baroclinic response with high pressure anomalies at the surface over the cold waters off
Newfoundland and low pressure anomalies off Irelandover warm SST conditions. The model’s low-level response over the ocean is thus opposite to that from
NCEP–NCAR. South of 308N, both model and observation exhibit a baroclinic profile in agreement with theclassical tropical/subtropical ocean–atmosphere relationship (Hoskins 1987). At higher latitudes (north of658N), the barotropic properties of the atmospheric flow
are recovered in the model.
A couple of hypotheses may explain NCEP–NCAR/
ARPEGE disagreements for late summer months at midlatitudes. The first one would refer to the nature of theexperimental design, which could artificially perturb the

low-level atmospheric response. It is well known thatthe so-called Atmospheric Model Intercomparison Project (AMIP) type simulations do not correctly capture
the surface ocean–atmosphere interactions as developedin Barsugli and Battisti (1998) and may affect the barotropic shape of the atmospheric modes that dominates
at midlatitudes. Such a mechanism might be seasonallydependant and could dominate in ARPEGE in summertime. A second hypothesis would be to consider themodel as realistic and to assume that it correctly responds to the extratropical SST anomalies. The anomalous low-level circulation in the model then would tend
to damp the surface ocean anomalies and would be consistent with the idea that the extratropical atmosphereis not responsible for the genesis of the extratropicalanomalous HS lobes. Such a hypothesis reinforces thetropical origin of the anomalous extratropical atmospheric circulation suggested previously. It has been verified that no significant changes are simulated in ARPEGE within the tropical band. In particular, the Atlantic
ITCZ is neither affected in intensity nor in position (notshown). In such a scenario, the switch between summer
baroclinic and fall quasi-barotropic model response
would also be considered as a realistic physical mechanism. The HS anomalies would then move from a passive role in summer to an active role in winter in theforcing of the extratropical atmosphere. In the rest of
the paper, we further document such a hypothesis andfocus our interest on the origin of the simulated baro-

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FIG. 8. Anomalous vertical profile for (a) zonal wind and (b) geopotential height for Sep–Nov months and averaged
over the 308–408N, 508–108W domain. Units are respectively (m s 21 ) and (m).

tropic response in fall. This will allow us to provide
some possible explanations about the observed link between the phase of the summer oceanic HS mode andthe early winter atmosphere.
c. Role of the climatological background in thesimulated atmospheric response
Figure 8 shows anomalous vertical profiles of zonal
wind and geopotential height in September, October, andNovember for the region (308–408N, 508–108W), whichcorresponds to the diffluence zone of the upper-tropospheric westerly wind jet and encompasses the maximum U200 westerly anomaly simulated by the model(Fig. 6d). While all three months exhibit positive vertical wind shear anomalies, the shear is amplified considerably in October and November compared to September (Fig. 8a). The vertical structure of the geopotential height response in the months with strong verticalwind shear is approximately equivalent barotropic
whereas it is baroclinic in the months with weak shear(Fig. 8b). Why do similar SST anomaly patterns leadto different vertical structures of atmospheric circulation
response in the presence of vertical wind shears?
Some studies suggest that anomalous transients andassociated feedbacks onto the mean flow play an important role in shaping the midlatitude atmospheric response to extratropical SST anomalies (Held et al. 1989;
Watanabe and Kimoto 1999). The relationship between
synoptic-scale eddies along the midlatitude storm trackand local vertical shear has been extensively documented in the literature (see, e.g., Hoskins and Valdez
1990). The model anomalous eddy activity estimatedby the variance of the bandpass (2.2–6 days) filtered
Z500 is presented in Fig. 9 as a function of latitude andfor the same longitudinal band used in Fig. 8, for themonths September, October, and November. The monthly U200 simulated anomalies are also shown. In Sep-

tember, although U200 westerly anomalies dominate
significantly between 308 and 508N, the storm activity(hereafter STA) is not significantly altered. In Octoberhowever, westerly U200 anomalies between 308 and
458N are accompanying by a significant enhancementin STA. In November, despite a slight 58 misfit, enhanced (reduced) STA coincides rather well in location
over the entire eastern basin with U200 anomalous acceleration (slackening) and is overall amplified. Thus,
it appears that increased vertical wind shear is accompanied by enhanced eddy activity changes from Octoberonward suggesting, all together, a modification of thebaroclinicity.
The alteration of the transients as well as their seasonality are presumably of central importance to explain
the nature of the mean atmospheric response (Peng etal. 1995). Climatological monthly mean tendencies in
U200 and STA computed as the difference between thegiven month and the previous one for the same region(308–408N, 508–108W) are shown in Fig. 9d. In September the mean STA is at its minimum amplitude,
while the upper-level climatological jet already starts
increasing over the selected area. It is hypothesized herethat the baroclinic instability criteria are not met yet for
September at this location. The jet might not be strong
enough to trigger changes in storminess, despite locallyenhanced vertical shear (Fig. 8a) interpreted as a direct
response to the anomalous meridional temperature gradient imposed at the surface. From October onward
U200 and STA are intensifying with maximum strengthening in November for the latter, one month later for
the former. This crude analysis tends to indicate that thedevelopment of synoptic storms in the model requires
a minimum speed of the jet. When such a threshold isovertaken, storms rapidly reach their maturity and do
not follow as closely the gradual increase of the upperlevel westerly circulation like in December. Due to amore favorable climatological background starting in

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3359

FIG. 9. Anomalous zonal wind at 200 hPa (U200) and storm track activity (STA) as a function of latitude andaveraged over the longitudinal 408–108W band for (a) Sep, (b) Oct, and (c) Nov. Units are m s 21 and m 2 . Thicksuperimposed lines exceed the 95% significance limit using T statistics. (d) Monthly changes for climatological U200
and STA, estimated as the difference between the mean value of the considered field for a given month and the one
for a month earlier.

October, it is suspected, in our case, that the presenceof local vertical shear, in relation to local prescribed
SST anomalies, is now able to affect the local baroclinicity and related storminess at midlatitudes. It is thussuspected that the action of the storm explains to a great
extent the barotropic nature of the mean atmospheric
response from October. This hypothesis relies on theproperties of the synoptic eddies to favor quasi-barotropic profiles as detailed, for instance, in Hoskins etal. (1983).
Beside its impact for explaining the nature of theatmospheric response, the alteration of the midlatitude
STA may provide a strong positive feedback in maintaining and intensifying the modified mean flow. It is
expected to affect the planetary-scale waves and to modify the stationary wave activity allowing for large-scale

anomalous circulation. The latter clearly appears in November in the model as detailed below and is clearly
amplified in December (not shown). The main storm
track is significantly reduced over the entire North Atlantic from the Labrador Sea to the British Isles, while
its southern edge is reinforced within a 308–458N latitudinal band (Fig. 10a). Such a pattern is consistent with
the negative phase of the NAO as described in Hurrell(2003). The simulated pattern bears some resemblancewith NCEP–NCAR changes estimated by November
lagged regression of the bandpass-filtered Z500 on the
AS SVD PC time series (Fig. 10c), especially east of
408W. ARPEGE tends, however, to underestimate thealteration of the synoptic eddies on the western side of
the basin (Newfoundland area) at midlatitudes and tooverestimate the storminess reduction farther north.

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FIG. 10. Linear ARPEGE model response in Nov for the (a) STA (m 2 ) and (b) for the Eliassen–Palm (EP) flux. The mean model zonal
wind at 200 hPa (U200) climatology is superimposed (m s 21 ; thick solid line) and contours start at 20 m s 21 , every 10 m s 21 . Regressionof NCEP–NCAR Nov (c) STA and (d) EP onto AS SST time series obtained from SVD. Contour intervals for STA are 0.5 m 2 and shaded
areas exceed the 95% significance limit using T statistics. The mean NCEP–NCAR U200 (thick solid line) is superimposed (d).

Such a discrepancy can be attributed to the mean biasof the model (position and strength of the storm track
and the jet) and the reader is invited to refer to Cassou
and Terray (2001) or Drévillon et al. (2003) for further
details.
Anomalous momentum fluxes associated with synoptic eddies are presented in Figs. 10b,d for the modeland NCEP–NCAR observations using the Eliassen–
Palm (EP) E vector following Trenberth’s (1986) definition. The zonal E component can be interpreted as thezonal stretching of the mean flow induced by the synoptic eddies and its meridional component as the dominant eddy forcing of the mean streamfunction evolution(Held et al. 1989). The divergence (convergence) of E
is indicative of mean flow acceleration (deceleration)
due to the presence of STA changes. A clear divergenceoccurs in the model on the southeastern side of the mean
climatological jet (around 308N and within 408–108W)
and is compensated by a strong convergence at its tailend off western Europe and at the entrance of the African subtropical jet. Such an anomaly is consistent withenhanced zonality of the basic flow and diminution of
the planetary wave activity (Doblas-Reyes et al. 2001).
It is also consistent with the forcing tendency due tosynoptic eddies to develop a negative NAO-type largescale pattern, as the strength of the Azores high (Icelandic low) is affected by the change in the main courseof the storms. Despite a clear 108 displacement to thenorth that can be related again to the model mean bias,
a similar pattern can be found in NCEP–NCAR. The
EP divergence (convergence) is maximum along thesouthern (northern) edge of the mean jet over the easternbasin (east of 408W).
4. Conclusions
a. Synthesis
Seasonal ocean–atmosphere interaction in the Atlantic has been investigated to obtain a better understandingof the statistical relationship between the summer SST
anomalies associated with the ‘‘Horseshoe’’ mode andthe NAO in the following early winter. A combined
observational and modeling approach was adopted to
examine the origin of the HS pattern and to provide
some physical hypotheses for the delayed HS forcing
on the next winter atmospheric circulation.
The results collectively suggest that HS can be primarily considered as an extratropical oceanic footprintof tropical convection changes in the Atlantic ITCZ via
atmospheric teleconnection mechanisms. When the climatological ITCZ is enhanced or displaced northward,
the atmospheric circulation is characterized by a cyclonic circulation off Newfoundland and diminished1 SEPTEMBER 2004

CASSOU ET AL.

trade winds in the subtropics. The midlatitude atmospheric alteration is speculated to be the surface signature of forced Rossby waves initiated within the Tropics. This hypothesis is confirmed from model experiments using CCM3, where diabatic heating anomalies
diagnosed from observations are imposed along the
ITCZ. Anomalous convection in the western tropicalbasin is responsible in CCM3 for the excitation of aforced wave train extending northeastward from the
Caribbian basin toward Scandinavia. Modifications in
trade wind intensity in the northern tropical Atlantic
occurs at the same time as a response to altered rainfall
over the Sahel. The model teleconnection pattern reproduces fairly well the observed anomalous low-levelatmospheric flow related to HS. The associated turbulent
and radiative fluxes at the sea surface would tend to
warm up the eastern and subtropical basins and cool theregion east of Newfoundland, as observed. This allows
us to conclude that the HS can be viewed as the local
imprint of tropically induced extratropical atmosphericchanges.
The late summer SST anomalies associated with theHS pattern were hypothesized to influence the atmospheric circulation over the North Atlantic and Europe
in late fall/early winter. To test this observational hypothesis, experiments with ARPEGE were performed
where the observed evolution of the HS mode is prescribed from August to November in the model SST
boundary conditions. The model response is characterized by a baroclinic structure through September withhigh (low) surface pressure above cold (warm) SST
anomalies, followed by an equivalent barotropic structure from October onward that resembles the NAO. The
baroclinic response can be interpreted as the atmospheric tendency to damp the extratropical SST anomalies
and may confirm the fact that the extratropical atmosphere by itself is not responsible for the HS. The barotropic response can be considered as the signature oflarge-scale atmospheric circulation in response to theextratropical SST anomalies. We have shown that anomalous vertical shear collocated with the anomalous meridional SST gradient is associated with changes instorm activity starting in October. The alteration of thestorm track is hypothesized to be of central importanceto initiate and sustain large-scale barotropic anomalies
through eddy–mean flow interaction. The model seems
to be more receptive to the SST anomalies when thestorminess and associated upper-level jet are well developed. This emphasizes the importance of the climatological background to explain the sign and the timing of the atmospheric response to the persistent HS
SST anomalies (see also Peng et al. 1995 but for late
winter).
b. Discussion
The analyses presented here from observations andfrom model experiments have insisted on the importance3361

of both summer tropical and extratropical conditions forpossibly influencing the next winter’s NAO. Tropicalatmospheric changes seem to initiate a bridge between
summer and early winter, and the resulting extratropical
SST anomalies seems to be of importance as they persist
enough to feed back to the atmosphere when the climatological background state is favorable. Therefore, it
is important to consider together, in a global perspective,
tropical and extratropical signals as they are not independent and contribute, by a cooperative and constructive interaction, to the variability of the winter NAO. It
would be interesting, in particular, to test the ocean–
atmosphere link presented in this paper from coupledmodel integrations.
In such a context, we can expand the present seasonal
summer-to-winter bridge and propose a winter-to-winterbridge to explain the year-to-year persistence of the
NAO and part of its low frequency. Let us now startwith the winter–early spring season (from January tolate April) and a given phase of the NAO. As describedin the introduction, the prevailing forcing of the atmosphere on the underlying ocean during the winter–
early spring months imprints a tripolelike structure of
SST anomalies, which tends to be maximum in spring(Cayan 1992). Anomalous trade wind intensity peaks atthat time (Hastenrath and Greischar 1993) and tends tobuild or reinforce the so-called oceanic interhemispheric
SST gradient in the deep Tropics. As detailed in Sutton
et al. (2000), the atmospheric response to the oceanic
cross-equatorial gradient is maximum for late spring/summer months (June–September) and is characterized
by a meridional displacement of the ITCZ toward thewarmer hemisphere. We have now established a link
between the summer tropical change and the previous
winter NAO conditions. Vimont et al. (2001) have diagnosed a similar connection, but for the Pacific Ocean.
Based on that link and on the present study on the relationship between summer SST conditions and the next
winter’s NAO, the same NAO phase as the previous
winter would be favored and a new cycle would startwith the reappearance of the same-sign winter North
Atlantic SST tripole. Summer months might thus appear
very important since, first, they are sensitive to the tropical SST anomalies forced to a great extent by the previous winter–spring NAO and, second, they prepare theextratropical oceanic HS anomalies to which the atmosphere is sensitive at the beginning of next winter.
This temporal bridge over summer might explain partof the redenning of the winter NAO spectrum.
How can this bridge or cycle be interrupted? We haveshown that the key vector sustaining the global mechanism is associated with the tropical convection over
the western part of the Atlantic basin and over the Sahel
to a lesser extent. A simple analysis of variance (ANOVA) shows that the internal atmospheric variability
dominates in this region despite being located within
the Tropics. It represents between 55% and 70% of thetotal variance depending on season, as estimated from

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JOURNAL OF CLIMATE

ARPEGE (Cassou and Terray 2001), and can therefore
disrupt the bridge. In addition to this internal component, El Niño–Southern Oscillation is one obvious candidate to alter the Atlantic cycle as it remotely affects
the strength and the position of the tropical convergencezones (Klein et al. 1999; Sutton et al. 2000; Giannini
et al. 2001; Alexander et al. 2002). The local interaction
between the land and the atmosphere over the SouthAmerican continent might also play a role (Geldney andValdez 2000), as well as the Indian Ocean variability
via its effect on the branch of the Walker circulation
over the African continent. The so-called ‘‘AtlanticNiño’’ (Zebiak 1993; Czaja and Frankignoul 2002) isalso expected to modify the convection pattern of thetropical Atlantic and ultimately affects the midlatitudes(Drévillon et al. 2003).
Finally, the decadal variability of the subtropical andsubpolar gyres (Sutton and Allen 1997) might modulate
the efficiency of the remote tropical influences in imprinting extratropical SST anomalies, depending on thephase of the low frequency oceanic mode. Above all,
we have to keep in mind that the internal dynamics ofthe extratropical North Atlantic atmosphere is by far
dominant, even in summer, and can also create its own
oceanic surface anomalies that may affect the following
winter NAO. A large amount of variance of the total
extratropical variability is controlled by internal atmospheric processes and the present study only explains
about 15%–20% of the signal (based on the ARPEGE
model response using ANOVA techniques, Cassou andTerray 2001).
Nevertheless, the lagged relationship between thesummer SST anomalies and the following winter atmospheric circulation anomalies might be exploited forpredictability purposes (Rodwell and Folland 2002).
The scenario presented here, however, may be too limited due to the linear constraint on the analyses. Following a nonlinear approach used in Cassou et al. (2004)
based on cluster analysis, it appears that the negative
phase of the NAO is indeed preferably linked to theprevious summer HS mode, whereas the positive phaseof the NAO is better associated with a different summer
SST structure more reminiscent of the SST tripole. It
would therefore be interesting to further investigate theasymmetry between the positive and negative phase of
the NAO and their relationship with previous oceanic
conditions.
Acknowledgments. This work is supported by the European Community via the 5th framework program
PREDICATE project under Contract EVK2-CT-199900020 and by NOAA under Grant NA06GP0394. The
authors want to thank A. Giannini, R. Sutton, J. Chiang,
and K. Trenberth for stimulating discussions. The firstauthor is grateful to E. Maisonnave, S. Valcke, and C.
Périgaud for their constant support and valuable critiques and to A. Phillips for help accessing and plotting
the data with the NCL software.

VOLUME 17
REFERENCES

Alexander, M. A., I. Bladé, M. Newman, J. R. Lanzante, N. C. Lau,
and J. D. Scott, 2002: The atmospheric bridge: The influence of
ENSO teleconnections on air–sea interaction over the global
oceans. J. Climate, 15, 2205–2231.
Barsugli, J., and D. S. Battisti, 1998: The basic effects of atmosphere–
ocean thermal coupling on midlatitude variability. J. Atmos. Sci.,
55, 477–493.
Battisti, D. S., U. S. Bhatt, and M. A. Alexander, 1995: A modelingstudy of the interannual variability in the wintertime North Atlantic Ocean. J. Climate, 8, 3067–3083.
Branstator, G., and S. E. Haupt, 1998: An empirical model of barotropic atmospheric dynamics and its response to tropical forcing. J. Climate, 11, 2645–2667.
Cassou, C., and L. Terray, 2001: Oceanic forcing of the wintertime
low-frequency atmospheric variability in the North Atlantic European sector: A study with the ARPEGE model. J. Climate, 14,
4266–4291.
——, ——, J. W. Hurrell, and C. Deser, 2004: North Atlantic winterclimate regimes: Spatial asymmetry, stationarity with time, andoceanic forcing. J. Climate, 17, 1055–1068.
Cayan, D. R., 1992: Latent and sensible heat flux anomalies over thenorthern oceans: Driving the sea surface temperature. J. Phys.
Oceanogr., 22, 859–881.
Czaja, A., and C. Frankignoul, 1999: Influence of the North Atlantic
SST on the atmospheric circulation. Geophys. Res. Lett., 26,
2969–2972.
——, and ——, 2002: Observed impact of Atlantic SST anomalieson the North Atlantic Oscillation. J. Climate, 15, 606–623.
de Coëtlogon, G., and C. Frankignoul, 2003: Persistence of wintersea surface temperature in the North Atlantic. J. Climate, 16,1364–1377.
Déqué, M., C. Dreveton, A. Braun, and D. Cariolle, 1994: The climate
version of Arpege/IFS: A contribution to the French community
climate modelling. Climate Dyn., 10, 249–266.
Deser, C., and M. S. Timlin, 1997: Atmosphere–ocean interaction on
weekly timescales in the North Atlantic and North Pacific. J.
Climate, 10, 393–408.
——, M. A. Alexander, and M. S. Timlin, 2003: On the persistenceof sea surface temperature anomalies in midlatitudes. J. Climate,
16, 57–72.
Doblas-Reyes, F. J., M. Déqué, F. Valero, and D. B. Stephenson, 1998:
North Atlantic wintertime intraseasonal variability and its sensitivity to GCM horizontal resolution. Tellus, 50A, 573–595.
——, M. A. Pastor, M. J. Casado, and M. Déqué, 2001: Wintertime
westward-travelling planetary scale perturbations over the Euro–
Atlantic region. Climate Dyn., 17, 811–824.
Drévillon, M., L. Terray, P. Rogel, and C. Cassou, 2001: Midlatitude
Atlantic SST influence on European winter climate variability
in the NCEP–NCAR Reanalysis. Climate Dyn., 18, 331–344.
——, C. Cassou, and L. Terray, 2003: Model study of the wintertime
atmospheric response to fall tropical Atlantic SST anomalies.
Quart. J. Roy. Meteor. Soc., 129, 2591–2611.
Fontaine, B., S. Janicot, and P. Roucou, 1999: Coupled ocean–atmosphere surface variability and its climate impacts in the tropical Atlantic region. Climate Dyn., 15, 451–473.
Frankignoul, C., and K. Hasselmann, 1977: Stochastic climate models. Part II: Application to sea surface temperature variability
and thermocline variability. Tellus, 29, 284–305.
——, A. Czaja, and B. L’Hévéder, 1998: Air–sea feedback in the
North Atlantic and surface boundary conditions for ocean models. J. Climate, 11, 2310–2324.
Gedney, N., and P. J. Valdez, 2000: The effect of Amazonian deforestation of the Northern Hemisphere circulation and climate.
Geophys. Res. Lett., 27, 3053–3056.
Giannini, A., M. A. Cane, and Y. Kushnir, 2001: Interdecadal changes
in the ENSO teleconnection to the Caribbean region and theNorth Atlantic Oscillation. J. Climate, 14, 2867–2879.
——, R. Saravanan, and P. Chang, 2003: Oceanic forcing of Sahel

1 SEPTEMBER 2004

CASSOU ET AL.

rainfall on interannual and interdecadal time scales. Science, 302,
1027–1030.
Hastenrath, S., and L. Greischar, 1993: Circulation mechanisms related to northeast Brazil rainfall anomalies. J. Geophys. Res.,98, 5093–5102.
Held, I. M., S. W. Lyons, and S. Nigam, 1989: Transients and theextratropical response to El Niño. J. Atmos. Sci., 46, 163–174.
Hoerling, M. P., J. W. Hurrell, and T. Xu, 2001: Tropical origins forNorth Atlantic climate change. Science, 292, 90–92.
Hoskins, B., 1987: Diagnostics of forced and free variability in theatmosphere. Atmospheric and Oceanic Variability, H. Cattle,
Ed., Royal Meteorological Society, 57–74.
——, and R. Pierce, 1983: Large Scale Dynamical Processes in the
Atmosphere. Academic Press, 397 pp.
——, and P. J. Valdez, 1990: On the existence of storm tracks. J.
Atmos. Sci., 47, 1854–1864.
——, I. James, and G. White, 1983: The shape, propagation andmean flow interaction of large scale weather systems. J. Atmos.
Sci., 40, 1595–1612.
Hulme, M., and M. G. New, 1997: Dependence of large-scale precipitation climatologies on temporal and spatial sampling. J. Climate, 10, 1099–1113.
Hurrell, J. W., 2003: An overview of the North Atlantic Oscillation.
Geophys. Monogr. Series, No. 134, Amer. Geophys. Union, 1–
22.
——, J. J. Hack, B. A. Boville, D. L. Williamson, and J. T. Kiehl,
1998: The dynamical simulation of the NCAR Community Climate Model version 3 (CCM3). J. Climate, 11, 1207–1236.
——, M. P. Hoerling, A. Phillips, and T. Xu, 2004: Twentieth century
North Atlantic climate change. Part I: Assessing determinism.
Climate Dyn., in press.
Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-Year Reanalysis Project. Bull. Amer. Meteor. Soc., 77, 437–471.
Kiehl, J. T., J. J. Hack, G. B. Bonan, B. A. Boville, B. P. Briegleb,
D. L. Williamson, and P. J. Rasch, 1996: Description of the
NCAR Community Climate Model (CCM3). NCAR Tech. Note
NCAR TN-4201STR, 152 pp. [Available from NCAR, Boulder,
CO 80307.]
Klein, S. A., B. J. Soden, and N. G. Lau, 1999: Remote sea surface
temperature variations during ENSO: Evidence for a tropicalatmospheric bridge. J. Climate, 12, 917–932.
Kushnir, Y., W. A. Robinson, I. Bladé, N. M. J. Hall, S. Peng, and
R. T. Sutton, 2002: Atmospheric GCM response to extratropical
SST anomalies: Synthesis and evaluation. J. Climate, 15, 2233–
2256.
Marshall, J., and Coauthors, 2001: Review: North Atlantic climate
variability: Phenomena, impacts and mechanisms. Int. J. Climatol., 21, 1863–1898.
Mehta, V. M., M. J. Suarez, J. V. Manganello, and T. Delworth, 2000:
Oceanic influence on the North Atlantic Oscillation and associated Northern Hemisphere climate variations, 1959–1993.
Geophys. Res. Lett., 27, 121–124.
Norris, J. R., Y. Zhuang, and J. W. Wallace, 1998: Role of low clouds
in summertime atmosphere–ocean interactions over the North
Pacific. J. Climate, 11, 2482–2490.
Okumura, Y., S. P. Xie, A. Numaguti, and Y. Tanimoto, 2001: TropicalAtlantic air–sea interaction and its influence on the NAO. Geophys. Res. Lett., 28, 1507–1510.
Peng, S., and J. S. Whitaker, 1999: Mechanisms determining theatmospheric response to midlatitude SST anomalies. J. Climate,12, 1393–1408.
——, and W. A. Robinson, 2001: Relationships between atmospheric

3363

internal variability and the responses to an extratropical SST
anomaly. J. Climate, 14, 2943–2959.
——, L. A. Mysak, H. Ritchie, J. Derome, and B. Dugas, 1995: The
difference between early and midwinter atmospheric responses
to sea surface temperature anomalies in the northwest Atlantic.
J. Climate, 8, 137–157.
——, W. A. Robinson, and S. Li, 2003: Mechanisms for the NAO
responses to the North Atlantic SST tripole. J. Climate, 16,
1987–2004.
Robinson, W. A., 2000: Review of WETS—The workshop on extratropical SST anomalies. Bull. Amer. Meteor. Soc., 81, 567–577.
Rodwell, M. J., and C. K. Folland, 2002: Atlantic air–sea interaction
and seasonal predictability. Quart. J. Roy. Meteor. Soc., 128,
1413–1443.
——, D. P. Rowell, and C. K. Folland, 1999: Oceanic forcing of thewintertime North Atlantic Oscillation and European climate. Nature, 398, 320–323.
Sardeshmukh, P. D., and B. J. Hoskins, 1988: The generation of global
rotational flow by steady idealized tropical divergence. J. Atmos.
Sci., 45, 1228–1251.
Stephenson, D. B., H. Wanner, S. Brø̈nnimann, and J. Luterbacher,2003: The history of scientific research on the North Atlantic
Oscillation. Geophys. Monogr., No. 134, Amer. Geophys. Union,
37–50.
Sutton, R. T., and M. R. Allen, 1997: Decadal variability of North
Atlantic sea surface temperature and climate. Nature, 388, 563–
567.
——, and D. L. R. Hodson, 2003: The influence of the ocean on
North Atlantic climate variability 1871–1999. J. Climate, 16,
3269–3313.
——, S. P. Jewson, and D. P. Rowell, 2000: The elements of climate
variability in the tropical Atlantic region. J. Climate, 13, 3261–
3284.
——, W. A. Norton, and S. P. Jewson, 2001: The North Atlantic
Oscillation—What role for the ocean? Atmos. Sci. Lett., 1, 89–
100.
Terray, L., and C. Cassou, 2002: Tropical Atlantic sea surface temperature forcing of the quasi-decadal climate variability over the
North Atlantic–Europe region. J. Climate, 15, 3170–3187.
Trenberth, K. E., 1986: An assessment of the impact of transient
eddies on the zonal flow during a blocking episode using localized Eliassen–Palm flux diagnostics. J. Atmos. Sci., 43, 2070–
2087.
Vimont, D. J., D. S. Battisti, and A. C. Hirst, 2001: Footprinting: A
seasonal link between the midlatitudes and tropics. Geophys.
Res. Lett., 28, 3923–3926.
Wallace, J. M., and D. Gutzler, 1981: Teleconnections in the geopotentiel fields during the Northern Hemisphere winter. Mon.
Wea. Rev., 109, 784–812.
Watanabe, M., and M. Kimoto, 1999: Tropical–extratropical connection in the Atlantic atmosphere ocean variability. Geophys. Res.
Lett., 26, 2247–2250.
——, and ——, 2000: Atmosphere–ocean thermal coupling in the
North Atlantic: A positive feedback. Quart. J. Roy. Meteor. Soc.,
126, 3343–3369.
Webster, P. J., 1982: Seasonality in the local and remote atmospheric
response to sea surface temperature anomalies. J. Atmos. Sci.,
39, 722–733.
Woodruff, S. D., H. F. Diaz, J. D. Elms, and S. J. Worley, 1998:
COADS Release 2 data and metadata enhancements for improvements of marine surface flux fields. Phys. Chem. Earth,23, 517–527.
Zebiak, S., 1993: Air–sea interaction in the equatorial Atlantic region.
J. Climate, 6, 1567–1586.