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The Effects of North Atlantic SST and Sea Ice Anomalies on the Winter Circulation in
CCM3. Part II: Direct and Indirect Components of the Response
CLARA DESER
National Center for Atmospheric Research,* Boulder, Colorado

GUDRUN MAGNUSDOTTIR
Department of Earth System Science, University of California, Irvine, Irvine, California

R. SARAVANAN AND ADAM PHILLIPS
National Center for Atmospheric Research,* Boulder, Colorado(Manuscript received 16 January 2003, in final form 6 October 2003)
ABSTRACTThe wintertime atmospheric circulation responses to observed patterns of North Atlantic sea surface temperature and sea ice cover trends in recent decades are studied by means of experiments with an atmospheric generalcirculation model. Here the relationship between the forced responses and the dominant pattern of internallygenerated atmospheric variability is focused on. The total response is partioned into a portion that projects onto
the leading mode of internal variability (the indirect response) and a portion that is the residual from thatprojection (the direct response). This empirical decomposition yields physically meaningful patterns whose
distinctive horizontal and vertical structures imply different governing mechanisms. The indirect response, which
dominates the total geopotential height response, is hemispheric in scale with resemblance to the North Atlantic
Oscillation or Northern Hemisphere annular mode, and equivalent barotropic in the vertical from the surface tothe tropopause. In contrast, the direct response is localized to the vicinity of the surface thermal anomaly (SSTor sea ice) and exhibits a baroclinic structure in the vertical, with a surface trough and upper-level ridge in thecase of a positive heating anomaly, consistent with theoretical models of the linear baroclinic response to
extratropical thermal forcing. Both components of the response scale linearly with respect to the amplitude of
the forcing but nonlinearly with respect to the polarity of the forcing. The deeper vertical penetration of anomalous
heating compared to cooling is suggested to play a role in the nonlinearity of the response to SST forcing.

1. Introduction
Sea surface temperature (SST) and sea ice cover in
the North Atlantic have undergone pronounced trends
in recent decades. Although driven in large part bychanges in atmospheric flow (cf. Deser et al. 2000; Seager et al. 2000), the trends in SST and sea ice may exert
a significant feedback upon the atmospheric circulation.
The nature of this feedback is the subject of the present
study and a companion study (Magnusdottir et al. 2004,
hereafter referred to as Part I). In this work, we examine
the wintertime atmospheric circulation responses to observed patterns of North Atlantic SST and sea ice cover*The National Center for Atmospheric Research is sponsored by
the National Science Foundation.
Corresponding author address: Dr. Clara Deser, Climate and Global Dynamics Division, NCAR, P.O. Box 3000, Boulder, CO 803073000.
E-mail: cdeser@ucar.edu

q 2004 American Meteorological Society

trends by means of experiments with an atmosphericgeneral circulation model, version 3 of the Community
Climate Model (CCM3). As shown in Part I, a notable
result of these experiments is the similarity of the response to trends in ice cover and to trends (with reversedsign) in SST, despite large differences in the spatial
pattern and amplitude of the two forcing parameters. In
addition, these responses resemble the model’s leading
mode of internal variability, the simulated equivalent of
the observed North Atlantic Oscillation or Northern
Hemisphere annular mode patterns (e.g., Wallace 2000).
Previous modeling studies have also found a correspondence between the patterns of internal variability
and the response to an imposed midlatitude heating perturbation (Peng and Robinson 2001; Hall et al. 2001)
or to increasing greenhouse gas concentrations (Kushneret al. 2001).
The purpose of Part II of this study is to further examine the relationship between the forced response totrends in SST and sea ice cover, and the dominant pat-

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FIG. 1. Jan (left) SST and (right) sea ice forcing used in the SST25 and ICE2 perturbation
experiments, respectively. In the SST panel, dark (light) shading denotes positive (negative)
anomalies relative to the control run; the contour interval is 1.5 K and the zero contour has been
omitted. In the sea ice panel, dark (light) shading denotes grid boxes where sea ice has been
removed (added) relative to the control run.

tern of internally generated atmospheric variability. Specifically, we partition the total geopotential height response into a portion that projects onto the leading modeof internal variability (the indirect response) and theresidual from that projection (the direct response). Thisempirical decomposition is shown to yield physicallymeaningful patterns whose distinctive horizontal andvertical structures imply different governing mechanisms. Kushner et al. (2001) also found such an approach insightful for understanding the zonally averaged
atmospheric circulation response in the Southern Hemisphere to increasing greenhouse gas concentrations ina coupled climate model.
It was noted in Part I that the total geopotential height
response to trends in SST and sea ice is nonlinear withrespect to the polarity of the forcing. Here we expand
upon this issue by examining the sensitivity of both thedirect and indirect components of the response to thesign of the imposed SST and sea ice anomaly trends.
We also investigate the vertical distribution of anomalous heating compared to cooling in these experiments,
a factor which may play a role in the nonlinearity of
the response.
The outline of the paper is as follows. Section 2 provides a brief overview of the experiments analyzed. Section 3 presents the decomposition of the total geopotential height response into direct and indirect components, and documents the vertical distribution of anomalous heating associated with the imposed SST and sea
ice anomaly trends. Section 4 discusses and summarizes
the results.
2. Experimental design
A complete description of the atmospheric generalcirculation model (CCM3) and experimental design isgiven in Part I. The two primary experiments analyzed
here are termed SST25 and ICE2. In SST25, the forcing is defined as the observed monthly SST anomaly
trend during 1954–94 over the North Atlantic, multiplied by a factor of 25 (i.e., the polarity of the observed
trend is reversed and its amplitude is magnified fivefold). In ICE2, the forcing is the observed monthly trend
in sea ice extent over the North Atlantic during 1958–
97, magnified by approximately a factor of 2. In eachcase, the forcing is applied as an anomaly upon the meanseasonal cycle. Each experiment, as well as a control
simulation forced with the climatological seasonal cycleof SST and sea ice cover, is integrated for a minimumof 61 yr; we analyze years 2–61 from each run.
Figure 1 shows the January SST and sea ice anomalies
in SST25 and ICE2, respectively; anomalies in the other winter months are similar (not shown). In SST25,
the primary forcing is a positive SST anomaly center
located in the subpolar gyre (maximum amplitude ;7K), with a weaker negative anomaly off the east coastof North America (maximum amplitude ;3 K). In ICE2,
the forcing consists of a reduction in sea ice cover in
the Greenland Sea and an extension of the ice edge inthe Labrador Sea. Note that a grid cell is either ice freeor 100% ice covered; there is no fractional ice concentration.
We shall also make use of additional experiments thathave been conducted to test the sensitivity of the response to the magnitude and polarity of the imposed
thermal forcing. For SST forcing, these are termed
SST15, SST12.5, and SST22.5; for sea ice forcing
these are ICE1, ICELAB, and ICEGRN. SST15 is identical to SST25 except for a sign reversal (e.g., the observed SST trend magnified by a factor of 5); SST12.5(SST22.5) is identical to SST15 (SST25) except thatthe observed trend is multiplied by a factor of 2.5

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FIG. 2. Decomposition of the (left) total Dec–Apr Z 500 responses in (top) SST25 and (bottom)
ICE2 into a component that projects onto the leading EOF of (middle) the control run and (right)
the residual from that projection. The contour interval is 10 m in all panels; positive (negative)
contours are solid (dashed) and the zero contour is the thin solid line. The locations of the SST
and sea ice anomalies are indicated by shading in the residual panels.

(22.5); ICE1 is identical to ICE2 except that the areal
extent of the observed sea ice trend has not been magnified; ICELAB (ICEGRN) is identical to ICE2 exceptthat only the Labrador (Greenland) Sea portion of thesea ice cover is altered.
3. Resultsa. Geopotential height response
The left-hand panels of Fig. 2 show the total geopotential height response at 500 hPa (Z 500 ) for the winterseason December–April for SST25 and ICE2, obtainedby subtracting the 60-yr mean of the control simulationfrom the 60-yr mean of each perturbation experiment(April is included in the definition of winter due to thestrong similarity between the March and April responses
in both experiments; not shown). Values exceeding approximately 10 m in absolute value are significant at
the 95% confidence level according to a local t test (notshown). Although the SST and sea ice forcings differ
in location and magnitude, the responses are remarkably
similar in both amplitude and spatial pattern. Both response patterns are hemispheric in nature (although theforcing is confined to the Atlantic sector), with positive
height anomalies at high latitudes and negative heightanomalies at midlatitudes of the Atlantic and Pacific.
Maximum height anomalies are ;80 m in the polar
region and ;2(20–40 m) in midlatitudes.
As noted in the introduction, there is an emerging
consensus that the internal variability in an atmosphericgeneral circulation model may play a strong role in shap-ing the pattern of the forced response. The middle panelsof Fig. 2 show the leading EOF of winter Z 500 from thecontrol run, which accounts for 33% of the variance(and is well separated from the second EOF) over the
Northern Hemisphere north of 308N. Cosine weighting
has been used in the covariance matrix to compute the
EOF. (The scaling of the EOF patterns, which differs
slightly in the two middle panels, is described shortly.)
The similarity between the leading EOF of the control
run and the response patterns in both experiments isevident: the pattern correlation is 0.70 for SST25 and0.84 for ICE2. The principal component time series forthe control run EOF is approximately normally distributed about a mean of zero (not shown but see Part I).
Although there is a high degree of similarity between
the leading Z 500 EOF in the control run and the Z 500
responses in the two experiments, there are some subtle
distinctions in the relative amplitude and location of thecenters of action. To examine these differences, we projected the response onto the EOF using linear leastsquares spatial regression, and then subtracted the projection (the EOF scaled by the regression coefficient:
middle panels of Fig. 2) from the response to obtain theresidual. We excluded the forcing region (908W–458E)
from the calculation of the spatial regression coefficientso as to optimally compare the features of the far-field
response that are common to the internal variability;
however, similar results are obtained when the full hemispheric domain is used (not shown). In both experiments, the total response is dominated by its projection
onto the internal variability; however, a well-defined

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FIG. 3. As in Fig. 2 but for different levels in the troposphere: (a) ICE2, (b) SST25, and (c)SST15.

residual pattern that is closely tied to the location of theforcing is also apparent (right-hand panels of Fig. 2).
In SST25, the residual response consists of an anomalous ridge over and slightly downstream of the warm
SST anomaly, with maximum amplitude ;60 m (10 m
K 21 ). In ICE2, the main feature of the residual response
is an anomalous ridge centered over the Greenland Sea
in the region of sea ice removal and extending over thepolar cap, with maximum amplitude ;50 m. The anomalous ridge response in the residual fields in both ex-periments is consistent with simple physical reasoning(e.g., a raising of geopotential height surfaces within
and above an anomalously warm column of air), lending
credence to our diagnostic technique and the notion thatthe total response may be partitioned into direct (e.g.,
residual) and indirect (e.g., projection onto the leading
mode of internal variability) components. We note thatneither residual pattern resembles any of the higherorder EOFs in the control run, suggesting that the residual is indeed a direct, forced response.

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FIG. 3. (Continued )The vertical structures of the total, indirect, and direct
responses are shown in Fig. 3. The indirect responses
were computed by regressing the geopotential heightanomalies (relative to the 60-yr mean) in the control
run at each level upon the leading 500-hPa principal
component time series in the control run, and then scaling those regression patterns (denoted Zreg ) by the spatial
regression coefficient between the total height response(experiment-minus-control) and Zreg at each level separately. In this way, we ensure that the indirect responses

shown in Fig. 3 represent the projections upon a single
mode of internal variability as represented by the leading PC time series at 500 hPa. However, using the leading EOF computed separately for each level in the control run in place of Zreg yields nearly identical results
due to the high temporal correlation (.0.96) between
the leading PC time series at each level with that at 500hPa (not shown).
For ICE2 (Fig. 3a), the vertical structure of the total
response pattern is approximately equivalent barotropic,

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FIG. 3. (Continued )with embedded small-scale structures evident at 1000hPa coincident with the location of the sea ice anomalies. The indirect response is also equivalent barotropic,
as expected and consistent with previous studies (e.g.,
Ting and Lau 1993). Note the absence of localized icerelated structures in the indirect response at 1000 hPa.
The direct response is baroclinic between 1000 and 850
hPa over the regions of anomalous sea ice cover; forexample, where sea ice has been removed in the Greenland Sea, the height anomalies are negative at 1000 hPa

and positive at 850 hPa, indicative of a local warmingof the boundary layer. Above the boundary layer, thedirect response consists of a primary center of above
normal heights over the Greenland Sea and spreading
to much of the Arctic. This anomalous ridge grows withheight from 850 to 500 hPa, with no further amplification between 500 and 300 hPa. A weaker anomalous
trough is also present over and downstream of the enhanced ice cover in the Labrador Sea.
For SST25 (Fig. 3b), the vertical structures of the1 MARCH 2004

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FIG. 4. As in Fig. 2 but for (top) SST22.5 and (bottom) ICE1.

total and indirect responses are approximately equivalent barotropic, similar to those for ICE2. The residual
response is baroclinic within the boundary layer directly
over the positive SST anomaly (i.e., an anomalous
trough at 1000 hPa and near-zero height anomalies at
850 hPa), also similar to the results for ICE2. The residual response in the free atmosphere consists of anequivalent barotropic anomalous ridge over and downstream of the warm SST anomaly. Note that this anomalous ridge amplifies considerably between 500 and 300
hPa, unlike the corresponding feature in ICE2, which
attains its maximum amplitude at 500 hPa.
We have also carried out the decomposition for
ICEGRN, ICELAB, and SST15. The total, indirect, anddirect responses for ICEGRN (not shown) are very similar to those for ICE2 (except for the lack of a local
boundary layer response over the Labrador Sea due tothe absence of sea ice forcing there), while the ICELAB
responses (not shown) are weak and not statisticallysignificant (except for the boundary layer response over
the Labrador Sea). These results indicate that it is theremoval of sea ice in the Greenland Sea rather than the
addition of sea ice in the Labrador Sea that is primarily
responsible for the circulation response in ICE2 (seealso Part I).
The total, indirect, and direct responses for SST15
are shown in Fig. 3c. While similar in spatial and vertical structure (albeit with opposite sign) to their SST25
counterparts, their magnitudes are considerably weaker.
For example, the amplitudes of the total and indirect
responses throughout the troposphere are approximately
a factor of 3 smaller for SST15 than SST25. In addition, the anomalous trough in the direct free-atmosphere response to SST15 exhibits less amplification

with height than the analogous ridge in SST25: themagnitude of the maximum height anomaly at 300 hPa
is ;30 m in SST15 compared to ;90 m in SST25.
The fact that our empirical technique for decomposing
the total response yields meaningful patterns even forthe small amplitude response in SST15 lends further
support to its credibility.
In summary, the total response to positive surface
temperature anomalies, either in the form of SST changes or reductions in sea ice cover, is dominated by thespatial (hemispheric) and vertical (approximately equivalent barotropic) structure of the leading mode of internal variability of the model’s wintertime atmosphericcirculation. The structure of this ‘‘indirect’’ componentof the response contrasts with that of the residual component, which is more locally tied to the region of forcing and is baroclinic within the boundary layer while
amplifying with height in the free atmosphere. The amplitude of the model response (total, direct, and indirect)
is much greater for a positive surface temperature perturbation than a negative one, both for SST and sea ice
forcing.
In contrast to the asymmetry in the strength of the
response with respect to the polarity of the forcing, themodel responds in a nearly linear fashion with respect
to the amplitude of the forcing. Figure 4 shows the total,
indirect, and direct Z 500 responses for SST22.5 andICE1 that may be compared with their doubled-forcing
counterparts SST25 and ICE2, respectively (recall Fig.
2). The spatial patterns of the responses are similar between the two SST experiments and between the two
sea ice experiments, and the amplitudes scale approximately linearly with the magnitude of the forcing. The
spatial regression coefficients between the Z 500 respons-

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FIG. 5. Vertically integrated heating rate anomalies for (top) the entire depth of the model
atmosphere (1000–10 hPa), (bottom) the boundary layer (1000–850 hPa), and (middle) the freetroposphere (750–250 hPa) for (left) SST15 and (right) SST25. The contour interval is 17.5 W
m 22 in all panels and the zero contour has been omitted.

es in SST22.5 and SST25, a measure of the relative
amplitudes of the response patterns are 0.50, 0.46, and0.49 for the total, indirect, and direct components, respectively, and those between ICE1 and ICE2 are 0.50,
0.54, and 0.38, respectively.
b. Nonlinearity of the response: The role of thevertical distribution of anomalous heating
One factor that may play a role in the stronger model
response to positive versus negative SST anomalies isthe nonlinear dependence of evaporation upon SST according to the Clausius–Clapeyron relation. Indeed, themagnitude of the latent heat flux anomaly averaged
within the 2 K (absolute value) contour of the main SST
anomaly center is larger by a factor of 1.4 in SST25
than SST15 (see maps in Part I figures). However, thisratio is not large enough to directly account for theapproximately three-fold increase in the amplitude of
the total response in SST25 compared to SST15.
Another factor that may be important is the vertical
distribution of the anomalous heating. Figure 5 shows
the anomalous heating rate fields integrated over theentire depth of the model atmosphere (1000–10 hPa),

the boundary layer (1000–850 hPa), and the free troposphere (750–250 hPa) for SST15 and 2SST15. In
both experiments, the anomalous heating integrated overthe depth of the atmosphere is localized to the regionsof anomalous SST. There is also anomalous heating overthe far eastern North Atlantic that is of opposite sign
to that farther west; we speculate this is due to thethermodynamic adjustment of the low-level southwesterly flow to the upstream SST perturbation, as in theexperiments of Hall et al. (2001). The magnitude of theanomalous heating in the free troposphere over the central North Atlantic is approximately 2.5 times larger in
SST25 than in SST15. The anomalous heating patterns
within the boundary layer are relatively linear by comparison.
Further detail on the vertical heating distribution is
given in Fig. 6, which shows cross sections of the anomalous heating rates in SST25 and SST15 averaged overthe latitude band (438–658N) as a function of longitudeand height. The top panel shows the total heating rateanomalies and the lower panels show the individualcomponents that make up the total: the sum of the radiative and vertical diffusion terms, condensational
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FIG. 6. Vertical cross sections for the lat band 438–658N of anomalous heating rates for(left) SST15 and (right) SST25 due to (middle top) deep convection, (middle bottom)
shallow convection, and (bottom) radiation plus vertical diffusion. (top) The sum of thelower three panels. The contour interval is 0.3 K day 21 in all panels; positive (negative)
values are indicated by solid (dashed) contours, and the zero contour has been omitted.

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FIG. 7. Vertical profiles of the total anomalous heating rate averaged
over the main SST anomaly center (438–658N, 508–308W) in SST15(dashed) and SST25 (solid). The SST15 profile has been multipliedby 21.

ing in total precipitation minus that in shallow convection (termed ‘‘deep convection’’ in the figure). It is clearthat the total heating rate differences in the free atmosphere (e.g., deeper penetration of anomalous heatingcompared to cooling) are almost entirely due to differences in the deep convective component that peaks inthe midtroposphere near 650 hPa. At this level, the magnitude of the anomalous heating due to deep convection
is approximately a factor of 3 larger for SST25 than
SST15. The anomalous heating rates due to shallow
convection and radiation plus diffusion are nearly linear
by comparison. Figure 6 also shows that the eastward
tilt with height apparent in the total heating rate is dueto the eastward displacement of deep convection relative
to shallow convection.
A summary of the vertical structures of anomalous
heating compared to cooling over the main SST anomaly
center in SST25 and SST15 is given in Fig. 7, whichshows area averages for the region (438–658N, 508–
308W; note that the sign has been inverted for SST15
for easier comparison with SST25). Although theanomalous heating within the boundary layer is similar
in magnitude for the two cases, it is considerably stronger and deeper in the free atmosphere for the positive
SST anomaly compared to the negative one. For example, the anomalous heating at 600 hPa is ;0.7 K
day 21 in SST25 compared to ;20.2 K day 21 in
SST15, and the level to which the anomalous heating
extends is ;300 hPa in SST25 compared to ;500 hPa
in SST15.
Given the stronger and deeper anomalous heating in

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SST25 compared to the anomalous cooling in SST15,
it is now evident why the direct component of the geopotential height response in SST25 amplifies withheight at a considerably greater rate than that in SST15(recall Fig. 3) because the thickness between two geopotential height surfaces must be proportional to thetemperature anomalies in the layer. The degree to which
the vertical profile of the anomalous heating or coolingaffects the strength of the indirect response cannot be
definitively answered without a more complete understanding of the dynamical mechanisms responsible forthe leading mode of internal variability in the model.
However, qualitatively we expect that a deeper heating
profile will project more strongly onto the equivalentbarotropic vertical structure of the leading internal modeof variability, thus exciting it more efficiently. In thiscontext, Hall et al. (2001) have examined the sensitivityof the response of a simplified AGCM to different idealized vertical profiles of midlatitude heating and findthat ‘‘the pattern of the global response is not very sensitive to the changes in the shallow heating profile, but
as the heating becomes deeper the response is stronger
both locally and in the equivalent barotropic teleconnections.’’ Indeed, we find that the indirect response isapproximately 3 times larger in magnitude for SST25
than SST15, a factor that is similar to the difference
in strength of the anomalous heating above the boundary
layer in the two experiments (recall Figs. 5–7).
Of course, the anomalous heating pattern itself may
be affected by the circulation response, in particular theindirect component. To examine this possibility, we decomposed the vertically integrated heating rate response
in SST25 into indirect and direct components by regression upon the leading Z 500 PC time series in thecontrol run [this time series is highly correlated (r 50.86) with the leading PC of the vertically integrated
heating field in the control run]. The projection onto theleading internal mode (Fig. 8) exhibits a north–south
dipole in the Atlantic, with enhanced heating along 408–
508N and diminished heating along 608–708N associatedwith a southward shift of the mean westerly jet (see also
Fig. 10a in Part I). However, this indirect component isapproximately a factor of 5 weaker than the total response, so that the total heating rate response is dominated by the direct component, unlike the case for geopotential height (recall Fig. 2). Thus, the indirect circulation response does not substantially impact the total
vertically integrated heating response. This result indicates that other processes besides heating, such as
eddy momentum fluxes, are important for the maintenance of the internally generated variability, consistent
with previous studies (cf. Ting and Lau 1993 and Fig.
14a in Part I). Analogous results are obtained for SST15
and ICE2 (not shown).
Why does the anomalous heating extend so muchhigher in the warm SST experiment compared to thecold one? The answer may be simply that cooling frombelow is an inherently stabilizing process while heating1 MARCH 2004

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FIG. 8. As in Fig. 2 but for the vertically integrated total anomalous heating rate in SST25.
The contour interval is 15 W m 22 in all panels (the values in the middle panel have been multipliedby 5) and the zero contour has been omitted.

from below is a destabilizing one, conducive to convective overturning and deeper vertical penetration. In
support of this notion, we show mean vertical temperature profiles averaged over the main SST anomaly center from the control, SST25, and SST15 experiments(Fig. 9). It is clear that the positive temperature perturbation is sufficiently buoyant relative to the control run
lapse rate that it can extend into the upper troposphere
while the negative one is mainly confined to the lower
troposphere.

FIG. 9. Long-term mean winter temperature profiles. The two coldest profiles are for the Greenland Sea region (748–858N, 208W–1008E)
from the control run (solid) and ICE2 (dashed). The three warmest
profiles are for the north-central Atlantic (438–658N, 508–308W) fromthe control run (solid), SST15 (thin dashed), and SST25 (thickdashed).

c. Sea ice versus SST forcing
In the previous section, we hypothesized that thestronger model response to positive versus negative SST
anomalies is mainly due to the deeper penetration of
anomalous heating compared to cooling. Does this argument apply to the sea ice experiments, and in particular, does it explain why the response to reduced ice
cover in the Greenland Sea (e.g., surface warming) is
so much larger than that to enhanced ice cover in the
Labrador Sea (e.g., surface cooling; not shown but see
Part I)? Figure 10 shows the anomalous heating rate
fields integrated over the entire depth of the model atmosphere, the boundary layer, and the free troposphere
for the ICE2 experiment. The anomalous heating rate
integrated over the depth of the atmosphere exhibits
large negative (positive) values over the enhanced (reduced) ice cover in the Labrador (Greenland) Sea, withweaker values of opposite sign directly downstream dueto the modification of the air mass as it encounters open
water (see Part I). Peak values associated with the ice
cover changes are ;2300 W m 22 over the Labrador
Sea and ;150 W m 22 over the Greenland Sea (the higher values in the Labrador Sea may be due to the stronger
cross-ice component of the near-surface winds that enhances the surface turbulent energy flux anomaly; notshown). These peak values are considerably larger thanthose in the SST anomaly experiments, although they
are restricted to a smaller area. In contrast to the resultsfor SST25, the anomalous heating over the Greenland
Sea is confined almost entirely to the lower troposphere.
We note that the shallowness of the anomalous heating
profile over the Greenland Sea is consistent with thevertical structure of the direct response (recall Fig. 3a)in that the anomalous ridge directly over the reduced
ice cover does not amplify above ;700 hPa.
The results shown in Fig. 10 raise two questions: 1)
why does the anomalous heating associated with reduced ice extent in the Greenland Sea not penetrate as
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the mean static stability in the region of the surface
temperature perturbation plays a decisive role in how
deep the anomalous heating can penetrate.
Given that the anomalous heating over the Greenland
Sea in the ICE2 experiment does not penetrate as deeply
as that over the North Atlantic in SST25, why are theindirect responses in free atmosphere comparable in thetwo experiments? Put another way, why is the leading
mode of internal variability excited so strongly in the
ICE2 (and ICEGRN) experiment? Although we do nothave a definitive answer to this question, we speculatethat the ice anomaly in the Greenland Sea may be located in a region to which the model’s internal variability is particularly sensitive. In this context, we notethat the Greenland Sea ice perturbation is broad in zonal
extent and located near the Arctic center of action of
the leading Z 500 EOF in the control simulation, perhaps
resulting in an efficient projection upon the spatial pattern of the dominant mode of internal variability. Another possible factor is the interaction of the flow in theforcing region with the upstream topographic barrier ofGreenland, which may result in a strong projection upon
the internal mode. Further work is planned to investigate
these issues.
4. Summary and further discussion
We have decomposed the simulated total geopotential
height responses to North Atlantic SST and sea ice coveranomalies into a portion that projects onto the leading
mode of internal variability (the indirect response) andthe residual from that projection (the direct response).
We suggest that this empirical decomposition yields
physically meaningful patterns with distinctive horizontal and vertical structures, as discussed below.
a. The direct response
FIG. 10. As in Fig. 5 but for ICE2.

the North Atlantic, and 2) if the anomalous heating overthe Greenland Sea is so shallow, why is the indirectcomponent of the geopotential height response so large(as large as the response in SST25)? The mean verticaltemperature profiles over the Greenland Sea (748–858N,
208W–1008E) in the control and ICE2 experiments,
shown in Fig. 9, provides some insight into the firstquestion. When sea ice is present (in the control run),the lower troposphere is nearly isothermal and consequently, very large positive surface temperature perturbations are needed to overcome the strong static stability. Indeed, a surface temperature anomaly of 17 K
(ICE2-minus-control) is only large enough to penetrate
to ;700 hPa. In contrast, a surface temperature perturbation of only 4 K in SST25 is sufficient to penetrate
to ;350 hPa due to the weaker stratification over the
ice-free North Atlantic (Fig. 9). Thus, we speculate that

The direct component of the response to a positive
North Atlantic SST anomaly (SST25) exhibits a baroclinic structure in the vertical, with a surface troughdirectly over the SST anomaly and an equivalent barotropic ridge in the free atmosphere over and slightly
downstream of the SST anomaly. This structure is qualitatively consistent with simple theoretical models of
the linear baroclinic response to extratropical thermal
forcing (see, e.g., the review by Kushnir et al. 2002).
A similar pattern of opposite sign is found for the direct
response to a negative SST anomaly (SST15), but theamplitude is weaker by a factor of 2–3. The weaker
response in SST15 than SST25 may be a result of the
reduced magnitude and vertical extent of the anomalous
cooling associated with a negative SST anomaly compared to anomalous heating associated with a positive
one. We postulate that the shallower vertical extent ofa negative heating anomaly compared to a positive one
is due to the stabilizing effect of surface cooling versus
the destabilizing effect of surface heating. Unlike the1 MARCH 2004

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asymmetry in the strength of the direct response to positive versus negative SST forcing, the direct response
is approximately linearly proportional to the magnitudeof the imposed SST anomaly.
The direct response to reduced sea ice cover in theGreenland Sea is analogous to the direct response in
SST25 except that the anomalous ridge in the free atmosphere does not penetrate as deeply (e.g., does notamplify with height above ;700 hPa). We suggest that
this is due to the more limited vertical extent of theanomalous heating resulting from the higher mean static
stability in the lower troposphere associated with thepresence of sea ice in the control run. As in the SST
forcing experiments, the direct response to anomalous
sea ice cover is nonlinear with respect to the polarityof the forcing and approximately linear with respect tothe amplitude of the forcing.
b. The indirect response
The indirect response, or the component of the response that projects onto the leading mode of internalvariability, dominates the total geopotential height response in both SST25 and ICE2. Like the direct response, the magnitude of the indirect response scales
linearly with the amplitude of the forcing but nonlinearly with the sign of the forcing. Unlike the direct
response, the horizontal structure of the indirect response is hemispheric in scale, resembling the North
Atlantic Oscillation or Northern Hemisphere annularmode, and the vertical structure is approximately equivalent barotropic from the surface to the tropopause.
Many outstanding issues regarding the indirect response remain to be investigated, including an understanding of the dynamical processes responsible for theleading internal mode of variability including essential
interactions between the transient and stationary eddies
and the mean flow, the factors that determine the polarityof the indirect response and the relative strengths of theindirect and direct responses, and the mechanisms responsible for the nonlinear behavior of the indirect response with respect to the sign of the forcing. Recently,
Peng et al. (2003) have proposed that nonlinear interactions between the anomalous heating and eddy-in-

duced components of the flow can lead to asymmetries
in the response to positive versus negative SST anomalies. Idealized experiments are planned to examine
these issues, including the sensitivity of the indirect response to the location and polarity of the forcing andto the vertical profile of the associated anomalous heating, as well as possible nonlinear interactions between
the anomalous heating and eddy forcing.
Acknowledgments. We wish to thank Dr. Michael Alexander for useful discussions during the course of thiswork and the three anonymous reviewers for helpfulcomments. We are also grateful to Dr. Christophe Cassou for his technical assistance. GM was supported by
NOAA OGP under Grant NA96GP0420.
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