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Simulation of the 1976/77 Climate Transition over the North Pacific: Sensitivity toTropical Forcing
CLARA DESER AND ADAM S. PHILLIPS
National Center for Atmospheric Research, Boulder, Colorado(Manuscript received 12 January 2006, in final form 10 April 2006)
ABSTRACTThis study examines the contribution of tropical sea surface temperature (SST) forcing to the 1976/77
climate transition of the winter atmospheric circulation over the North Pacific using a combined observational and modeling approach. The National Center for Atmospheric Research (NCAR) Community
Atmospheric Model version 3 (CAM3) simulates approximately 75% of the observed 4-hPa deepening of
the wintertime Aleutian low from 1950–76 to 1977–2000 when forced with the observed evolution of tropical
SSTs in a 10-member ensemble average. This response is driven by precipitation increases over the western
half of the equatorial Pacific Ocean. In contrast, the NCAR Community Climate Model version 3 (CCM3),
the predecessor to CAM3, simulates no significant change in the strength of the Aleutian low when forced
with the same tropical SSTs in a 12-member ensemble average. The lack of response in CCM3 is traced toan erroneously large precipitation increase over the tropical Indian Ocean whose dynamical impact is to
weaken the Aleutian low; this, when combined with the response to rainfall increases over the western andcentral equatorial Pacific, results in near-zero net change in the strength of the Aleutian low. The observed
distribution of tropical precipitation anomalies associated with the 1976/77 transition, estimated from a
combination of direct measurements at land stations and indirect information from surface marine cloudiness and wind divergence fields, supports the models’ simulated rainfall increases over the western half of
the Pacific but not the magnitude of CCM3’s rainfall increase over the Indian Ocean.

1. Introduction
The wintertime atmospheric circulation over the
North Pacific has undergone a substantial long-term
intensification since the late 1940s (Nitta and Yamada
1989; Trenberth 1990; Trenberth and Hurrell 1994;
Mantua et al. 1997; Minobe 1997; Zhang et al. 1997;
Deser et al. 2004; among others). This intensification,
manifest as a deepening of the climatological Aleutian
low pressure system, is responsible for widespread climatic changes over North America (Cayan et al. 2001;
Dettinger et al. 2000) and physical and biologicalchanges within the North Pacific Ocean (Miller et al.
1994; Deser et al. 1996; Mantua et al. 1997; Deser et al.
1999; Miller et al. 1998). Because of its relatively rapid
onset in the late 1970s, this long-term circulation variation has been termed the “1976/77 climate transition”(Nitta and Yamada 1989; Trenberth 1990; Miller et al.

Corresponding author address: Dr. Clara Deser, Climate andGlobal Dynamics Division, NCAR, P.O. Box 3000, Boulder, CO
80307-3000.
E-mail: cdeser@ucar.edu© 2006 American Meteorological Society

JCLI3963

1994; Graham et al. 1994). A similar long-term deepening of the Aleutian low occurred during the first halfof the twentieth century, followed by a strengthening(Minobe 1997; Zhang et al. 1997; Deser et al. 2004).
The current view of the origin of the North Pacific1976/77 climate transition is that it was initiated bychanges in tropical sea surface temperatures (SSTs)
over the Pacific and Indian Ocean sectors (Minobe
1997; Graham et al. 1994; Zhang et al. 1997; Deser et al.
2004). Since 1977, tropical SSTs have increased by approximately 0.4°C over the Indian Ocean and far western Pacific and by approximately 0.6°C over the eastern
Pacific relative to the period 1950–76 (Fig. 1). This pattern of SST change is similar to that which occurs during the warm phase of the El Nino–Southern Oscillation (ENSO) phenomenon, although the Indian Ocean
SST increases are relatively more prominent and theeastern Pacific SST changes are not narrowly confinedto the equatorial upwelling zone (Garreaud and Battisti
1999; Deser et al. 2004). The proposed mechanism forthe 1976/77 climate transition over the North Pacific issimilar to that which occurs on interannual time scales
in association with ENSO, namely, that tropical Indo-

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FIG. 1. Epoch difference map of observed winter (December–February) SST, obtained by subtracting
the period 1950–76 from the period 1977–2000, based on the HadISST dataset (Rayner et al. 2003). The
contour interval is 0.2°C, negative contours are dashed, and the zero contour is omitted; values ⬎0.2°C
are shaded.

Pacific SST anomalies give rise to precipitation anomalies whose associated diabatic heating forces an atmospheric Rossby wave train to the North Pacific (Horeland Wallace 1981) that is subsequently modified by interactions with transient eddies along the North Pacificstorm track (Trenberth et al. 1998; Barsugli andSardeshmukh 2002). As emphasized by DeWeaver andNigam (2004) with regard to ENSO events, it is the full
two-dimensional distribution of tropical rainfall anomalies (both positive and negative) that ultimately determines the pattern and amplitude of the extratropicalatmospheric response.
The hypothesis of tropical SST forcing of the 1976/77
climate transition over the North Pacific has beentested to some degree in experiments with atmosphericgeneral circulation models (GCMs) (Chen et al. 1992;
Graham 1994; Graham et al. 1994). The results of theseearly studies generally support the notion that tropical
SST forcing plays an important role in the deepening of
the Aleutian low since 1977. However, these early investigations contain shortcomings, including the relatively coarse model resolution, limited period of simulation, idealized SST forcing configuration, and singlemember integrations. In this study, we make use of amultimember ensemble of integrations with three stateof-the-art atmospheric GCMs forced by the observed
history of SSTs over the tropical oceans (20°S–20°N)
during the period 1950–2000 to more comprehensively
examine the contribution of tropical forcing to the
1976/77 climate transition over the North Pacific.
One difficulty in assessing the role of the Tropics isthe lack of adequate rainfall data needed to characterize the full spatial pattern of the tropical precipitation
anomalies associated with the 1976/77 transition. Be-fore the advent of satellite rainfall measurements in the
mid-1970s, our knowledge of rainfall variability was
limited to land stations, leaving vast areas of the tropical oceans unsampled. Recently, studies have demonstrated the utility of ship-based observations of cloud
cover and surface wind divergence (see Deser et al.
2004 and Norris 2005) as proxy measures of rainfall
over the tropical oceans. We adopt this novel approachto estimate the full spatial pattern of the tropical rainfall changes associated with the 1976/77 climate transition. This approach allows us to assess the realism of
the precipitation responses in the atmospheric GCMs
forced with observed tropical SSTs, experiments we use
to evaluate the role of tropical forcing in the intensification of the Aleutian low since 1950.
The paper is structured as follows. Section 2 describes the observational datasets, atmospheric GCM
experiments, and analysis methods. Section 3 describes
the results, and section 4 provides a summary and discussion.

2. Data, methods, and model integrations
a. Data
The following monthly averaged observational
datasets, available from http://www.cgd.ucar.edu/cas/
guide/, have been used in this study; we have analyzed
the period of record 1950–2000 unless stated otherwise.
1) Sea level pressure (SLP) from an updated version of
Trenberth and Paolino (1980). This archive is based
upon daily historical analyses for the Northern
Hemisphere poleward of 20°N on a 5° latitude ⫻ 5°
longitude grid, updated with data from the National

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Centers for Environmental Prediction–National
Center for Atmospheric Research (NCEP–NCAR)
40-Yr Reanalysis Project (Kalnay et al. 1996).
2) 500-hPa geopotential height from the NCEP–
NCAR reanalysis on a 2.5° latitude ⫻ 2.5° longitude
grid (Kalnay et al. 1996).
3) Land station precipitation from the University of
East Anglia’s Climate Research Unit (Hulme et al.
1998) on a 2.5° latitude ⫻ 3.75° longitude grid, andfrom Dai et al. (1997) on a 2.5° latitude ⫻ 2.5° longitude grid. The last year available is 1998 from
Hulme et al. and 1996 from Dai et al.
4) Total cloud amount and surface wind from the International Comprehensive Ocean–Atmosphere
Data Set (ICOADS; Woodruff et al. 1987; Worley etal. 2005; Woodruff et al. 2005) on a 2° latitude ⫻ 2°
longitude grid.
5) SST from the Hadley Centre Sea-Ice and Sea Surface Temperature (HadISST) (Rayner et al. 2003)
on a 1° latitude ⫻ 1° longitude grid.

b. Methods
We formed monthly anomalies by subtracting thelong-term monthly means from each calendar month.
Winter (December–February) means were computedfrom the monthly anomalies. We use simple differences
between means for the periods 1950–76 and 1977–2000
to characterize the changes during the second half ofthe twentieth century; however, linear trends based on
the full period of record (1950–2000) give nearly identical results (not shown). For the observational archives
that contain substantial amounts of missing data [e.g.,
station precipitation fields of Hulme et al. (1998) andDai et al. (1997), and the ICOADS], a minimum of 25%of the winters in each period were required to have data
at a given grid box; otherwise, the period mean for thatgrid box was set to missing. To improve the readabilityof the ICOADS cloudiness and surface wind divergence epoch difference maps, linear interpolation(across gaps not exceeding 5 points in longitude and 3
points in latitude) and weighted binomial smoothing (9points in longitude and 3 points in latitude) were applied. The statistical significance of the epoch difference values was assessed by means of Student’s t test
taking into account serial correlation (Zwiers and vonStorch 1995).
Changes in observing practice and instrumentation
have caused spurious increases in scalar wind speed andcloudiness since 1950 from the ICOADS (Ward andHoskins 1996; Norris 1999; Wu and Xie 2003; Norris2005). Following the recommendations of Norris(2005), we account for these spurious trends by removing the tropical (30°N–30°S) mean cloud amount fromVOLUME 19

each grid box in each winter, and by removing a 4.7%
per decade linear trend from the surface zonal and meridional wind components. These “corrections” are notcritical to the results shown in this study (i.e., we havecompared results with and without the corrections andfind only minor differences between them).

c. Model integrations
Integrations with three state-of-the-art atmosphericgeneral circulation models forced by observed monthly
SSTs over the tropical oceans (20°N–20°S) during
1950–2000 from the dataset of Hurrell et al. (2005,
manuscript submitted to J. Climate) are analyzed in thisstudy. These consist of a 12-member ensemble of the
National Center for Atmospheric Research Community Climate Model version 3 (CCM3) at T42 horizontal
resolution (equivalent to 2.8° latitude ⫻ 2.8° longitude),a five-member ensemble of Community Atmosphere
Model Version 3 (CAM3) at T42 resolution, and a fivemember ensemble of CAM3 at T85 resolution (equivalent to 1.4° latitude ⫻ 1.4° longitude). Details of thephysical and numerical methods used in CCM3 andCAM3 are provided in Kiehl et al. (1998) and Collins etal. (2006), respectively. The members of the ensemble
for each model differ only in their initial conditions.
Atmospheric greenhouse gas concentrations are kept
fixed throughout the integrations at their 1990 levels.
We shall refer to these integrations as “TOGA” (Tropical Ocean–Global Atmosphere) simulations following
convention. Each model simulation was forced by theobserved history of monthly SSTs in the Tropics (20°N–
20°S) and climatological monthly mean SSTs polewardof 30°; a linear transition was used between 20° and 30°.
In addition to the model experiments describedabove, we have also analyzed the output of a 20member ensemble of CCM3 integrations forced withthe observed linear SST trends over the tropical Indian
Ocean conducted by Hoerling et al. (2004); the reader
is referred to that paper for details.

3. Results
The change in winter SLP from 1950–76 to 1977–2000
over the North Pacific and adjacent continents is shownin Fig. 2 (left) from observations and the three different
TOGA simulations: CCM3, CAM3 (T42), and CAM3
(T85). The model results are based upon the ensemble
mean response of 5 (12) TOGA simulations of CAM3
(CCM3). In nature, there is a significant lowering of
SLP over the North Pacific during 1977–2000 relative to
1950–1976, with maximum amplitude ⬃4.5 hPa near thesouthern tip of the Aleutian Islands. Both CAM3 ensembles simulate a substantial portion of the observed

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FIG. 2. Epoch differences of winter (December–February) (left) sea level pressure (SLP; left) and (right) 500-hPa geopotential height(Z500), obtained by subtracting the period 1950–76 from the period 1977–2000, from observations and three atmospheric generalcirculation models forced with observed time-varying tropical sea surface temperatures (referred to as “TOGA” simulations): CAM3
at T85 horizontal resolution, CAM3 at T42 horizontal resolution, and CCM3 at T42 horizontal resolution. The contour interval for SLP(Z500) is 1 hPa (10 m), the zero contours are omitted, negative contours are dashed, and the shading indicates epoch difference values
that are statistically significant at the 95% level based on a Student’s t test.

signal, with maximum amplitude ⬃3.5 hPa in the T85
version compared to ⬃3 hPa in the T42 version; these
responses are significant at the 95% level. In contrast to
CAM3, the CCM3 TOGA ensemble exhibits near-zero
SLP change over the North Pacific and adjacent continents. The 500-hPa geopotential height epoch differ-ences (Fig. 2, right) exhibit significant negative anomalies over the North Pacific and positive anomalies
downstream over western Canada in both nature andthe CAM3 simulations (as with SLP, the CAM3 responses are weaker than observed, with maximum
negative anomalies ⬃50 m in observations compared to6174

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FIG. 3. Histogram of the area-averaged SLP epoch difference
response over the North Pacific (30°–65°N, 140°E–160°W) basedon the (top) 10 CAM3 TOGA integrations (T42 and T85 combined) and the (bottom) 12 CCM3 TOGA integrations. The
dashed vertical line indicates the observed value (⫺2.24 hPa).

⬃35 m in CAM3 T85 and ⬃30 m in CAM3 T42). In
contrast to CAM3 and similar to the results for SLP, theCCM3 TOGA ensemble exhibits near-zero heightchanges over the North Pacific and adjacent continents.
An indication of the spread of the epoch difference
response among the individual ensemble members isgiven in Fig. 3, which shows a histogram of the areaaveraged SLP response over the domain 30°–65°N,
140°E–160°W (also known as the “North Pacific index”
after Trenberth and Hurrell 1994) for the 10 CAM3
integrations (T42 and T85 combined; top panel) andthe 12 CCM3 integrations (bottom panel). Nine of the
10 CAM3 integrations exhibit negative values, with amean of ⫺1.92 hPa, which is close to the observed valueof ⫺2.24 hPa. In contrast, 8 of the 12 CCM3 integrations exhibit positive values, and none of the CCM3

VOLUME 19

integrations exhibits a value ⬍⫺1 hPa; the mean value
for the CCM3 ensemble is 0.25 hPa.
Why is CCM3 so different from CAM3 in its epoch
difference atmospheric circulation response to tropical
SST forcing? To gain insight into this question, we compare the simulated patterns of tropical precipitation
change that force the atmospheric teleconnection to theNorth Pacific (Fig. 4). The tropical rainfall patterns arevery similar in the two CAM3 simulations. Both exhibit
the largest rainfall changes over the western two-thirdsof the tropical Pacific Ocean, with positive anomalies(maximum values ⬃2.5–3.5 mm day⫺1) along the equator and negative anomalies directly to the north. Negative anomalies also occur over northern South America
and the adjacent Atlantic Ocean. Rainfall changes overthe Indian Ocean are much weaker than those over the
Pacific and are generally positive. The amplitudes of
the Pacific rainfall changes are ⬃25% larger in the T85
version of the model compared to the T42 version, afactor that may contribute to the stronger SLP responseover the North Pacific in T85 than T42.
The CCM3 tropical rainfall epoch difference simulation resembles CAM3 over the Pacific and Atlantic sectors; however, it differs greatly over the Indian Ocean
where large positive values (as large as those over theequatorial Pacific) occur. CCM3 also exhibits negative
rainfall changes over the far western Pacific near 15°N
extending from the South China Sea to east of the Philippines, while CAM3 is weakly positive in this region.
We hypothesize that the differences between the
CAM3 and CCM3 rainfall distributions over the Indian
Ocean and far western Pacific may contribute to thedifferences in atmospheric response over the North Pacific. In particular, several modeling studies have shownthat positive rainfall anomalies over the Indian Ocean
west of ⬃110°E and negative rainfall anomalies near
the Philippines force positive SLP and geopotential
height anomalies over the North Pacific, opposite to the
response to positive rainfall anomalies over the westernand central equatorial Pacific (Branstator 1985; Ting
and Sardeshmukh 1993; Barsugli and Sardeshmukh
2002; G. Branstator 2005, personal communication).
Thus, we conjecture that in CCM3 the positive rainfall
anomalies over the Indian Ocean (and negative rainfallchanges near the Philippines) may be canceling out theeffects of the positive rainfall anomalies over the equatorial Pacific in terms of their impact upon North Pacific SLP.
Before testing this hypothesis, we evaluate the realism of the simulated tropical rainfall epoch differencesby comparing them to the available observational archives. As discussed in the introduction, the only direct
measurements of rainfall before the satellite era are1 DECEMBER 2006

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FIG. 4. Epoch difference maps, obtained by subtracting the period 1950–76 from theperiod 1977–2000, of simulated winter precipitation from the three atmospheric generalcirculation models forced with observed time-varying tropical sea surface temperatures(referred to as “TOGA” simulations): CAM3 at T85 horizontal resolution, CAM3 at T42
horizontal resolution, and CCM3 at T42 horizontal resolution. The color scale is given at
the bottom of the figure in units of mm day⫺1.

those at land stations: as a result, there is limited coverage over the oceans. However, indirect information
on rainfall changes over the tropical oceans may beobtained from ship-based observations of cloud coverand surface wind divergence, measures closely associated with rainfall in the Tropics. Figure 5 compares
epoch differences of land station rainfall from thedataset of Hulme et al. [1998; similar results are obtained using Dai et al. (1997); not shown] with those of
surface wind divergence and total cloud amount fromICOADS. The land station data show rainfall increases
during 1977–2000 relative to 1950–76 over the central
and western equatorial Pacific (maximum values ⬃2.5–
3.5 mm day⫺1), flanked by decreases to the south andwest. Precipitation increases are also found over Javaand northwestern Australia and decreases over Malaysia, Borneo, and the Philippines. Data are sparse overthe eastern Pacific, Indian, and Atlantic Oceans.
Comparison of the marine surface wind divergenceand cloudiness epoch difference fields (Fig. 5, bottom)

to the land station precipitation field (Fig. 5, top) shows
generally good correspondence in regions of overlapping coverage, giving confidence to the use of thesemeasures as qualitative indicators of precipitation. In
addition, the surface wind divergence and cloudinessepoch difference fields exhibit close spatial agreement
and physical consistency, with areas of enhanced (diminished) cloudiness generally corresponding to regions of anomalous surface wind convergence (divergence): this agreement between two entirely independent sets of measurements strongly supports theveracity of their signals. The largest amplitude changes
occur over the central equatorial Pacific, with enhanced
convergence and cloudiness during 1977–2000 compared to 1950–76. [Prominent cloudiness changes in thisregion are also observed for the earlier North Pacificclimate transitions in the 1920s and 1940s (Deser et al.
2004).] Weaker signals of enhanced convergence andcloudiness are found over the Arabian Sea and Bay of
Bengal, the western half of the Indian Ocean south of

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FIG. 5. Epoch difference maps, obtained by subtracting the period 1950–76 from the period 1977–2000, of observed winter precipitation from the gridded land station dataset of Hulme et al. (1998), and of observed winter surface marine cloud amount (color shading)
and wind divergence (contours) from the ICOADS dataset. The color scale is given at the bottom of the figure for both precipitation(in units of mm day⫺1) and cloud amount (in units of oktas). The contour interval for surface wind divergence is 1⫻10⫺7 s⫺1, negative
contours (indicative of anomalous convergence) are dashed, and the zero contour is darkened.

the equator (albeit mixed with pockets of anomalous
surface wind divergence), and the equatorial and northwestern Atlantic. Diminished cloudiness and anomalous surface wind divergence cover much of the remaining tropical oceans, especially the Indian Ocean between the equator and 10°N, the eastern Indian Ocean
south of the equator, the western equatorial and southwestern tropical Pacific, and in the central Pacific along⬃10°N and 160°–110°W. The regions of opposing
cloudiness and surface wind divergence signals west ofSouth and North America, where diminished cloud
cover is associated with enhanced surface wind convergence, are also physically consistent as these are areas
dominated by low clouds and subsidence (Klein andHartmann 1993). In this regime, cloud cover diminishes
when subsidence (and surface divergence) declines, allowing drier air to be entrained into the atmosphericboundary layer from above the low-level inversion. We
note that these regions are largely nonprecipitating.
How do the simulated rainfall epoch difference maps
compare with observations? In general, the primary signal of positive precipitation anomalies over the equa-

torial Pacific west of ⬃110°W simulated in all threemodels (Fig. 4) is verified in the cloudiness, divergence,and station precipitation datasets (Fig. 5), although
they extend farther west in the models compared tonature. There is some observational evidence for thesimulated negative rainfall anomalies directly to thenorth of the equator over the Pacific, although thesimulated amplitudes are generally larger than observed, especially north and east of New Guinea. All
three models underestimate the rainfall changes along⬃20°S in the western Pacific, although CCM3 is somewhat better than CAM3 in this regard. The negative
rainfall anomalies over northern South America are
well simulated in all the models. While it is impossible
to verify the simulated rainfall changes over the Indian
Ocean with direct measurements due to lack of data,
the consistent signals in cloudiness and surface winddivergence over this region suggest that the large precipitation increases over the Indian Ocean simulated in
CCM3 are erroneous.
As discussed above, we wish to test the hypothesisthat the different North Pacific SLP responses in CCM3

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and CAM3 to tropical SST forcing associated with the
1976/77 transition are due to differences in their tropical rainfall responses. Figure 6a shows the difference
between the CCM3 and CAM3 (T42) precipitation and
SLP epoch difference maps. As expected from Fig. 4,
there are large, widespread positive rainfall differences
over the Indian Ocean and negative precipitation differences north of New Guinea. The positive rainfall
differences a few degrees south of the equator in thewestern half of the Pacific reflect that the positive rainfall epoch differences in CCM3 extend slightly farther
south than those in CAM3. The corresponding SLP
differences over the North Pacific are positive, withmaximum values exceeding 2.5 hPa.
To test whether the tropical precipitation differences
are responsible for the North Pacific SLP differences in
the CCM3 and CAM3 TOGA experiments, we haveforced CCM3 with the tropical (20°N–20°S) diabatic
heating epoch difference field from the ensemble meanof the CAM3 (T42) TOGA runs, and also with thatfrom the CCM3 TOGA runs, and differenced the two
experiments. Each experiment consists of adding thethree-dimensional diabatic heating epoch difference
anomalies to the monthly diabatic heating field from along control run with fixed monthly mean climatological SSTs, and integrating the model for 11 yr (the last 10
yr of integration are used in our analysis). Because themodel is free to adjust to the imposed diabatic heating
anomalies, the equilibrium diabatic heating anomalies
may differ from those initially imposed. Figure 6bshows the differences between the precipitation and
SLP responses in the two CCM3 experiments (imposed
CCM3 heating minus imposed CAM3 heating), scaled
by 0.38. This scaling factor was obtained by taking theratio of the area average of the absolute value of the
precipitation anomalies over the tropical Indian andwestern Pacific Oceans (20°N–20°S, 45°–175°E) between Fig. 6a and the CCM3 heating experiment differences (imposed CCM3 heating minus imposed
CAM3 heating). The scaling factor reflects that themodel strongly amplifies the initial diabatic heating
forcing (a similar scaling factor of 0.4 is obtained if thearea averaging is computed over the entire tropicalIndo-Pacific). While there is not a perfect match between the precipitation changes in Figs. 6a and 6b,
there is overall agreement, with positive precipitation
anomalies over the Indian Ocean and negative values
over the Maritime Continent region; positive values
also occur in the eastern half of the Pacific along ⬃5°N.
The corresponding SLP response (Fig. 6b, left) closely
resembles the SLP difference between CCM3 andCAM3 (Fig. 6a, left), with positive changes over the
North Pacific (maximum values exceeding 2.5 hPa) and

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negative changes over northeastern Siberia. The resultsof these experiments lend credence to the notion that
disparities between the tropical precipitation epoch difference responses in CCM3 and CAM3 are responsiblefor differences in their North Pacific SLP responses. A
caveat to this conclusion is that the precipitation differences between CCM3 and CAM3 (Fig. 6a) are not
identical to those simulated by the idealized CCM3 experiments (Fig. 6b): in particular, the former exhibits
weaker precipitation anomalies over the western Indian
Ocean and stronger rainfall changes north of New
Guinea than the latter.
To highlight the role of SSTs over the tropical Indian
Ocean in forcing biases in the CCM3 tropical precipitation epoch difference response, Fig. 6c shows themean precipitation and SLP responses from a 20member ensemble of CCM3 forced by the observed
linear trends in SSTs over the tropical Indian Ocean(25°N–25°S, 40°–110°E) from 1950 to 1999 (characterizing the SST forcing by means of linear trends or epoch
differences yields nearly identical patterns; not shown).
This experiment was conducted by Hoerling et al.
(2004) and made available to us by the authors. As withthe earlier CCM3 experiments, we have scaled the responses by the ratio (0.56) of the area average of theabsolute value of the precipitation changes over thedomain 20°N–20°S, 45°–175°E between Fig. 6a and theIndian Ocean SST forcing experiment. The precipitation response closely resembles that in Fig. 6b over theeastern hemisphere, indicating that positive Indian
Ocean SST anomalies force both rainfall increases locally and rainfall decreases remotely over Indonesia
and South Africa. In addition, the similarity of the SLP
response over the North Pacific to that in Fig. 6b indicates that a substantial portion (⬃75%) is due to tropical precipitation changes in the eastern hemisphere associated with increased SSTs over the tropical Indian
Ocean.

4. Discussion and conclusions
We have examined the contribution of tropical SST
forcing to the 1976/77 climate transition of the winteratmospheric circulation over the North Pacific using a
combined observational and modeling approach. We
found that CAM3 (at both T42 and T85 resolutions)
simulates ⬃75% of the observed 4 hPa deepening of
the wintertime Aleutian low from 1950–76 to 1977–
2000 when forced with the observed evolution of tropical SSTs, while, surprisingly, CCM3 (at T42 resolution)
simulates no significant change in the strength of the
Aleutian low. We traced the different atmospheric circulation responses in the two models to disparities in6178

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FIG. 6. (a) Difference between CCM3 and CAM3 (T42) epoch differences of winter (left) SLP and (right) precipitation. (b) Winter(left) SLP and (right) precipitation from CCM3 forced with its own tropical diabatic heating epoch difference field minus CCM3 forced
with the tropical diabatic heating epoch difference field from CAM3 (T42); see text for details. (c) Winter (left) SLP and (right)
precipitation responses in CCM3 forced with the observed linear SST trend from 1950 to 1999 over the tropical Indian Ocean (from
Hoerling et al. 2004). The contour interval for SLP is 0.5 hPa, negative contours are dashed, and the zero contour is omitted. The color
scale for precipitation is given below the right column of panels in units of mm day⫺1.

their tropical precipitation responses. While both models exhibit similar epoch rainfall increases over thewestern half of the equatorial Pacific, their rainfall increases over the tropical Indian Ocean differ greatly inmagnitude, with changes as large as those over the Pacific in CCM3 and much weaker anomalies in CAM3.
Our estimate of the actual tropical precipitation
changes associated with the 1976/77 transition, basedon a combination of direct measurements at land stations and indirect information from surface marine
cloudiness and wind divergence fields, compares morefavorably to the CAM3 simulation than to CCM3.
Based on these results, we hypothesized that CCM3’serroneously large rainfall increases over the Indian
Ocean are dynamically consistent with the inability ofthe model to simulate the observed deepening of the
Aleutian low associated with the 1976/77 transition.
Further experiments aimed at testing this hypothesis
confirmed that increased precipitation over the tropicalIndian Ocean in CCM3 leads to reduced rainfall overIndonesia and the far western Pacific, which together
drive a positive SLP response over the North Pacificthat serves to cancel the negative SLP response due toincreased rainfall over the equatorial Pacific.
We have emphasized the role of tropical SST forcing(via increased rainfall over the western half of the equatorial Pacific) of the North Pacific 1976/77 climate transition. Other factors may also contribute, such as interactions between the tropically forced atmospheric circulation response over the North Pacific and theunderlying oceanic mixed layer, as well as air–sea coupling within the tropical Indian and western Pacific“warm pool” regions (Alexander et al. 2002). However,the direct forcing of the 1976/77 transition in the December–February atmospheric circulation over the
North Pacific by local SST changes is small according to
a comparison between the TOGA ensemble and an
ensemble integration forced by observed SSTs over the
global oceans (not shown).
Not only do our observational results based on direct

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and indirect measurements suggest there was no large
epoch increase in rainfall over the Indian Ocean associated with the 1976/77 climate transition, but our modeling results indicate such a precipitation change would
be dynamically inconsistent with the observed epoch
strengthening of the North Pacific winter atmosphericcirculation. Further research is needed to reconcile ourfindings with those of Hoerling et al. (2004) and Hurrell
et al. (2004) who attribute nearly half of the recent
upward trend in the winter (January–March) North Atlantic Oscillation from 1950 to 1999 to increased precipitation over the tropical Indian Ocean in response towarming of the underlying sea surface. Regardless, thedemonstrated sensitivity of the global extratropical atmospheric circulation to tropical precipitation changes
underscores the importance of continuing high-quality,
temporally homogeneous precipitation measurements
over the tropical oceans.
Acknowledgments. We acknowledge many helpful
discussions with Grant Branstator during the course ofthis study. We thank Michael Alexander, Grant Branstator, and Jim Hurrell for helpful comments on an
earlier draft of the manuscript, and the official reviewers for their suggestions and comments. We are gratefulto Andrew Mai and Bruce Briegleb for conductingsome of the model integrations used in this study. This
work was supported in part by a grant from NOAA’s
Office of Global Programs. The figures were producedwith the NCAR Command Language (NCL) software
package.
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