15 JANUARY 2014

WETTSTEIN AND DESER

527

Internal Variability in Projections of Twenty-First-Century Arctic Sea Ice Loss:
Role of the Large-Scale Atmospheric Circulation
JUSTIN J. WETTSTEIN AND CLARA DESER
National Center for Atmospheric Research,* Boulder, Colorado(Manuscript received 27 November 2012, in final form 7 June 2013)
ABSTRACTInternal variability in twenty-first-century summer Arctic sea ice loss and its relationship to the large-scaleatmospheric circulation is investigated in a 39-member Community Climate System Model, version 3 (CCSM3)
ensemble for the period 2000–61. Each member is subject to an identical greenhouse gas emissions scenario anddiffers only in the atmospheric model component’s initial condition.
September Arctic sea ice extent trends during 2020–59 range from 22.0 3 106 to 25.7 3 106 km2 across the 39
ensemble members, indicating a substantial role for internal variability in future Arctic sea ice loss projections.
A similar nearly threefold range (from 27.0 3 103 to 219 3 103 km3) is found for summer sea ice volume trends.
Higher rates of summer Arctic sea ice loss in CCSM3 are associated with enhanced transpolar drift andFram Strait ice export driven by surface wind and sea level pressure patterns. Over the Arctic, the covarying
atmospheric circulation patterns resemble the so-called Arctic dipole, with maximum amplitude between
April and July. Outside the Arctic, an atmospheric Rossby wave train over the Pacific sector is associated withinternal ice loss variability. Interannual covariability patterns between sea ice and atmospheric circulation are
similar to those based on trends, suggesting that similar processes govern internal variability over a broad
range of time scales. Interannual patterns of CCSM3 ice–atmosphere covariability compare well with those innature and in the newer CCSM4 version of the model, lending confidence to the results. Atmospheric teleconnection patterns in CCSM3 suggest that the tropical Pacific modulates Arctic sea ice variability via theaforementioned Rossby wave train. Large ensembles with other coupled models are needed to corroborate
these CCSM3-based findings.

1. Introduction
September Arctic sea ice extent has decreased byslightly more than 2.5 3 106 km2 (;10% decade21) since
1979, when consistent satellite observations were initiated (Fetterer et al. 2002, with updates; Serreze et al.
2007; Comiso et al. 2008; Deser and Teng 2008). Comprehensive Arctic sea ice thickness measurements are
more difficult to retrieve, but available measurements(Rothrock et al. 1999) and surrogate measures such as
the fraction of multiyear ice (Nghiem et al. 2006) and ice
age (Maslanik et al. 2007) indicate ice thickness and* The National Center for Atmospheric Research is sponsoredby the National Science Foundation.

Corresponding author address: Justin J. Wettstein, College ofEarth, Ocean, and Atmospheric Sciences, School of Public Policy,
Oregon State University, 104 CEOAS Administration Building,Corvallis, OR 97331.
E-mail: justinw@coas.oregonstate.edu
DOI: 10.1175/JCLI-D-12-00839.1
Ó 2014 American Meteorological Society

volume losses are also occurring. Most coupled climate
models forced with increasing greenhouse gas (GHG)
concentrations simulate reductions in Arctic summer sea
ice over the late twentieth century, a trend that accelerates over the twenty-first century (Stroeve et al. 2007,2012). Such models commonly forecast an essentially icefree (,1 3 106 km2) summer Arctic by mid to latetwenty-first century (e.g., Meehl et al. 2006; Stroeve et al.
2007; Kay et al. 2011; Meehl et al. 2012; Vavrus et al.
2012; Bitz et al. 2012; Stroeve et al. 2012). The magnitude
and rate of projected twenty-first-century Arctic sea ice
loss is subject to considerable uncertainty, however(Stroeve et al. 2007, 2012).
Uncertainty regarding the magnitude and rate of Arctic
sea ice loss stems from 1) different assumptions regarding
future external (GHG) forcing, 2) intermodel spread
resulting from different model physics and parameterizations, and 3) internally generated variability (e.g.,
Tebaldi and Knutti 2007; Hawkins and Sutton 2009; Deseret al. 2012b). By comparing observations to simulated
trends in a 6-member ensemble of the Community

528

JOURNAL OF CLIMATE

Climate System Model, version 4 (CCSM4), Kay et al.
(2011) estimated that approximately half (56%) of theobserved rate of 1979–2005 September Arctic sea ice
extent decline was externally forced, with the remainder
due to internal climate variability. A similar fraction(60%) of the observed 1979–2011 sea ice loss was attributed to external forcing by Stroeve et al. (2012) basedon a comparison with the multimodel ensemble mean ofsimulations from phase 5 of the Climate Model Intercomparison Project (CMIP5).
The sensitivity of Arctic sea ice loss projections to theamount of external forcing was evaluated in CMIP3 andCMIP5 by subjecting each model to different levels of
GHG increases. Significant differences in ice loss do not
emerge until approximately the middle of the twentyfirst century [e.g., Figs. 12 and 19, respectively, in Meehl
et al. (2006, 2012)] when atmospheric GHG concentrations associated with the different GHG emissions scenarios begin to diverge substantially [e.g., Fig. 5 in Mosset al. (2010)].
The intermodel spread in projected Arctic summer
sea ice decline is roughly comparable in magnitude
to the multimodel-mean estimate of greenhouse gas–
forced sea ice loss over at least the next several decades
in the CMIP3 and CMIP5 archives [e.g., Figs. 1 and 2,
respectively, in Stroeve et al. (2007, 2012)]. Intermodel
spread in ice loss is often accompanied by differences in
the intensity of local surface energy budget feedbacks(DeWeaver et al. 2008; Bitz 2008; Winton 2008; Boeet al. 2010; Kay et al. 2012), ocean heat transports(Holland and Bitz 2003; Mahlstein and Knutti 2011; Kayet al. 2012; Bitz et al. 2012), or atmospheric heat transports (Winton 2008). DeWeaver and Bitz (2006) foundsubstantially different climatological-mean surface wind
and ice advection patterns resulting in substantiallydifferent climatological-mean Arctic sea ice amounts
and distributions in two multicentury control simulations of the Community Climate System Model, version 3(CCSM3), when two different spatial resolutions wereprescribed for the component Community Atmosphere
Model, version 3 (CAM3). A strong model-dependent
relationship is also evident between the amount of Arctic
sea ice in the late twentieth century and the amount at theend of the twenty-first century in ensembles from both
CMIP3 [e.g., Fig. 1 in Stroeve et al. (2007)] and CMIP5[e.g., Fig. 2 in Stroeve et al. (2012)].
In contrast to the first two sources of uncertainty
outlined above, the role of internal variability in twentyfirst-century Arctic sea ice loss has received substantiallyless attention. Holland et al. (2008) noted that internalvariability plays an important role in CCSM3 simulations of abrupt losses in September sea ice extent. Kay
et al. (2011) and Jahn et al. (2012) described highly

VOLUME 27

variable Arctic sea ice loss trends in a small (6-member)set of CCSM4 simulations of the late twentieth century.
Kay et al. (2011) also noted that internal variability influences both the magnitude and sign of projected future
Arctic sea ice trends well into the twenty-first century.
The purpose of this study is to comprehensively evaluate
the uncertainty in future Arctic sea ice loss associatedwith internally generated climate variability, using asingle coupled climate model and a much larger (39member) set of simulations than in previous studies.
We also expand on earlier work by investigating themechanisms that control future internally generated
Arctic sea ice loss.
Observed interannual and decadal-scale variability in
Arctic sea ice extent has been related to variability inatmospheric circulation patterns that can be characterized as either regional (Vinje 2001; Tsukernik et al.
2010) or hemispheric in scale (Rigor et al. 2002; Rigor
and Wallace 2004; Belchansky et al. 2005; Overland andWang 2005; Wu et al. 2006; Ogi and Wallace 2007;
L’Heureux et al. 2008; Wang et al. 2009; Overland andWang 2010; Ogi et al. 2010; Screen et al. 2011; Stroeve
et al. 2011; Ogi and Wallace 2012). Smedsrud et al.
(2011) linked changes in Fram Strait sea ice export toanomalous Arctic wind patterns and pan-Arctic sea ice
variability. Observed sea ice variability has also been
related to anomalies in the thermodynamic exchangeof energy across the ice–ocean–atmosphere interface(Francis and Hunter 2007; Perovich et al. 2007) and toanomalies in ocean heat transport (Polyakov et al. 2005;
Shimada et al. 2006). We note that changes in the surface
energy budget and ocean heat transport may in turn beassociated with variations in the atmospheric circulation. In the studies that relate ice loss to large-scale atmospheric variability, there are divergent views on
the spatial structure and seasonality of the covariant
circulation anomalies. Covariant central Arctic surfacepressure patterns that resemble the Arctic Oscillation(AO; Thompson and Wallace 1998) are described bymany studies (Rigor et al. 2002; Rigor and Wallace 2004;
Belchansky et al. 2005; Ogi and Wallace 2007; Ogi et al.
2010; Stroeve et al. 2011; Ogi and Wallace 2012), whereas
patterns more consistent with the dipole anomaly (Wuet al. 2006) and Arctic dipole (AD; Overland and Wang2010) are identified in others (Overland and Wang 2005;
Wu et al. 2006; Wang et al. 2009; Overland and Wang
2010; Screen et al. 2011). L’Heureux et al. (2008) noted
an unusually strong midtropospheric anticyclone over thewestern Arctic and a Pacific/North American (PNA)-likepattern extending into the midlatitudes during themonths preceding the previous record low September sea
ice extent in 2007. Even though the studies above preferentially cluster around either AO- or AD-like patterns

15 JANUARY 2014

WETTSTEIN AND DESER

over the central Arctic, individual studies within eachgroup diverge substantially on the seasonality of the ice–
atmosphere relationships identified.
In this study, we investigate uncertainty in Arctic sea
ice loss projections over the coming decades due to internal variability in the climate system, with a particularfocus on the role of the atmospheric circulation. We
use a unique 39-member CCSM3 ensemble of climatechange simulations for the 2000–61 period. This set ofsimulations has previously been used to assess climate
predictability (Branstator and Teng 2010; Teng andBranstator 2011) and to compare anthropogenicallyforced and internally generated contributions to futureclimate change signals in temperature, precipitation,
and atmospheric circulation (Deser et al. 2012a,b; Hu
et al. 2012). We also make use of smaller twenty-firstcentury climate change simulation ensembles with thenewer version of the model, CCSM4. In addition, we
analyze the relatively short observational record to assess the realism of the physical relationships between
Arctic sea ice and atmospheric circulation variability on
interannual time scales as simulated by the models. The
model simulations, observational datasets, and analysis
methods are described in section 2. Results are presented in section 3 and summarized in section 4. A discussion concludes the paper.

2. Data and methods
The primary dataset used in this study is a unique ensemble of 39 simulations with the CCSM3, a fully coupled
non-flux-adjusted model with atmosphere, ocean, land,and sea ice components. Various aspects of the CCSM3
and its twenty-first-century response to several Special
Report on Emissions Scenarios (SRES) forcing trajectories are described in Collins et al. (2006) and Meehl
et al. (2006), respectively. As described in Deser et al.
(2012b), each member of this large CCSM3 ensemble issubject to the same SRES A1B external forcing scenario
during the twenty-first century and identical initialconditions in the ocean, sea ice, and land component
models obtained from 1 January 2000 at the end of aseparate CCSM3 twentieth-century simulation. The 39
different ensemble members were generated by initializing CCSM3’s atmospheric component model (CAM3)
with atmospheric conditions obtained from differentdays surrounding 1 January 2000 at the end of thetwentieth-century simulation. The 39-member ensemble
uses CAM3 at a T42 horizontal resolution (triangular
spectral truncation at a total wavenumber of 42; equivalent to 2.88 resolution in latitude and longitude) and 26
levels in the vertical. With the exception of an excluded
40th ensemble member that was found to have some529

corrupted output fields, it is the same ensemble used byDeser et al. (2012a,b) and a superset of the one used by Hu
et al. (2012). More information on this unique large
CCSM3 ensemble, especially regarding climate predictability, is available in Branstator and Teng (2010) and
Teng and Branstator (2011).
Smaller (seven or fewer member) ensembles fromCCSM4 that incorporate the newly released Community
Atmosphere Model, version 4 (CAM4), at a nominal
horizontal resolution of 18 are used to assess the robustness of CCSM3 results. Key CCSM4 features and its
response to various representative concentration pathway (RCP) external forcing scenarios are described in
Gent et al. (2011) and Meehl et al. (2012), respectively.
Satellite-based sea ice observations and an atmospheric reanalysis product are used to assess the applicability of the model results to nature. September Arctic
sea ice extent for 1979–2010 is obtained digitally fromthe National Snow and Ice Data Center (NSIDC; Fetterer
et al. 2002, with updates). We calculated monthly-mean
ice velocities for 1980–2006 from daily data retrieved by
the National Aeronautic and Space Administration’s(NASA) Polar Pathfinder satellite (Fowler 2003, with
updates). Monthly-mean sea level pressure (SLP) and
1000-hPa winds for 1980–2010 are obtained from theEuropean Center for Medium-Range Weather Forecasts(ECMWF) Interim Re-Analysis (ERA-Interim) product(Simmons et al. 2007, with updates).
We use two complementary indices of summer sea
ice amount: September extent and July–November(JASON) volume. The two metrics emphasize different
aspects of sea ice variability. In particular, the September
extent time series contains relatively more interannual
variability compared to the more slowly evolving JASON
volume record (not shown). The JASON averaging period is chosen based on the high (.0.6) month-to-monthautocorrelation of high-pass filtered volume time series
within this season; similar results are also obtained,
however, using September volume (not shown). Ice extent is defined as the total area (km2) of Arctic grid cellswith at least 15% sea ice concentration. In the models asin observations, Arctic sea ice is consistently least extensive during September and the largest ice extent loss
trends also occur during September. Summation calculations for both September ice extent and JASON ice
volume are performed using the native irregular grid of
the component sea ice model to avoid substantial errors
associated with spatial smoothing.
The T42 resolution CCSM3 simulates sea ice that is
too thick and extensive during the late twentieth century
relative to observations (DeWeaver and Bitz 2006).
Because of this, we begin our analysis in the year 2020
when the simulated sea ice is more realistic and we530

JOURNAL OF CLIMATE

examine trends over the 40-yr period 2020–59. As CCSM4
simulates a more realistic sea ice distribution than CCSM3(Jahn et al. 2012), we analyze 2010–49 output from thatmodel by applying similar considerations to those described above for CCSM3. We will discuss the influence ofdifferent climatological-mean sea ice distributions in
section 3 when we compare results from CCSM3 withthose from observations and CCSM4.
A few terms are defined here to save space and avoid
confusion. Extent (EXT) refers to simulated September
Arctic sea ice extent and volume (VOL) refers to simulated JASON Arctic sea ice volume during the 40-yr intervals described above for CCSM3 and CCSM4. Whether
we use linear 40-yr trends or interannual variability, theassociated extent and volume indices have been consistently multiplied by 21 so that larger (smaller) values indicate greater (lesser) ice loss. Geopotential heights at
1000 and 500 hPa are denoted as Z1000 and Z500, respectively. Zonal wind at 200 hPa is denoted as U200 and nearsurface reference air temperature is denoted as Tref.
This paper focuses on internal variability in summer
Arctic sea ice loss by analyzing the spread of trends
across the ensemble. The terms ensemble trend regression
and ensemble trend correlation will be shorthand for regressions onto and correlations with the 39 standardized
2020–59 EXT or VOL linear trend values obtained fromthe CCSM3 ensemble. These ensemble trend regressions
and ensemble trend correlations therefore describe relationships between internal variability in the trends ofvarious fields and the internal variability in sea ice loss
trends. All results based on linear trends are similar to
those based on epoch differences (not shown).
Empirical orthogonal function (EOF)—principal
component (PC) analysis is used to characterize theleading patterns of Z1000, Z500, and Tref trend variability
across the CCSM3 ensemble. Unlike conventional EOF
and PC analysis applied to the time domain, we apply
the method to the ensemble member trend domain (seealso Deser et al. 2012b). That is, the set of 39 different
2020–59 trend maps are subjected to EOF and PC
analysis. We refer to the resulting spatial patterns as
EOFs and the 39 weights for each spatial pattern as PCs.
Note that by construction, the ensemble-mean trends
are removed in the EOF and PC analysis. All EOF andregression patterns are displayed in physical units perstandard deviation of the reference index (e.g., PC,
EXT, or VOL).

3. Results
This section is structured as follows. First, we assess
the magnitude of the spread in projected Arctic sea ice
loss across the 39-member CCSM3 ensemble, spread

VOLUME 27

that is by definition attributable to internal variability(section 3a). Then we explore the role of the atmospheric circulation in driving internal variability in sea
ice loss trends, paying particular attention to their spatial pattern, seasonal dependence, and imprint upon ice
advection (section 3b). To help validate the model results, we compare the relationships between atmospheric circulation and Arctic sea ice variability on
interannual time scales in CCSM3, CCSM4, and observations (section 3c). Finally, we examine the global
signature of the atmospheric circulation patterns thatdrive sea ice loss trends and explore to what extent they
represent dominant structures of variability in their own
right, independent of their relationship to Arctic sea ice(section 3d).

a. The importance of internal variability on
projections of twenty-first-century Arctic sea
ice loss
The importance of internal variability on projectionsof twenty-first-century Arctic sea ice loss is demonstrated
in Fig. 1. Averaged across the 39 members, CCSM3 exhibits an ensemble-mean ice loss trend of 4.1 3 106 km2 in
EXT and 13 3 103 km3 in VOL over the 2020–59 period.
These ensemble-mean trends provide an estimate of the
GHG-forced signal in Arctic sea ice loss. However, individual ensemble members can differ markedly from theensemble mean because of the superposition of internallygenerated and externally forced trends. In fact, the EXTand VOL trends range by nearly a factor of 3 (from22.0 3 106 to 25.7 3 106 km2 in EXT and from 27.0 3103 to 219 3 103 km3 in VOL) across the ensemble, indicating the substantial contribution of internal variability to future Arctic sea ice loss (Fig. 1, top). The
2020–59 trend values have a standard deviation (range)of 0.83 3 106 km2 (3.7 3 106 km2) and 2.7 3 103 km3(12 3 103 km3) in EXT and VOL, respectively. Notethat each ensemble member is a plausible simulation of
the future trajectory of Arctic sea ice loss according to
CCSM3.
Figure 1 (top) illustrates the time series of EXT and
VOL for the ensemble members with the smallest andlargest trends. Not only do the 2020–59 trends differ
markedly (see histograms), but large interannual anddecadal-scale excursions about the ensemble-mean trend
are evident (note the higher proportion of low-frequency
variability in the VOL time series compared to the EXTtime series). Internal variability therefore has a discernible influence over a wide range of time scales. There islittle evidence in these time series for the developmentof a strong internal positive feedback whereby a chancenegative or positive departure from the ensemble-mean
trend either persists over the long term or amplifies as time

15 JANUARY 2014

WETTSTEIN AND DESER

FIG. 1. (top) Time series of the ensemble mean (black line), ensemble range (gray shading), and the individual
ensemble members with the smallest (blue lines) or largest (red lines) 2020–59 loss in Arctic (top left) September sea
ice EXT and (top right) July–November (JASON) sea ice VOL. Histograms to the right of the top panels display the
2020–59 ice loss trend distribution on the same vertical axis as the time series. Two small green dots indicate theensemble-mean ice loss trends in EXT and VOL. (middle),(bottom) Maps show the 2020–59 (left) ice concentrationor (right) thickness trends for the individual ensemble members with the largest and smallest trends in ice EXT or
VOL loss, respectively. The largest (smallest) trend in September ice EXT loss of 5.7 3 106 km2 (2.0 3 106 km2) is
experienced by ensemble member 9 (13), as indicated. The largest (smallest) trend in JASON ice VOL loss of 19 3
103 km3 (7.0 3 103 km3) is experienced by ensemble member 6 (13), as indicated. A scale for the 2020–59 trend maps
in September ice concentration (%) and JASON ice thickness (m) is indicated alongside the colorbar.

531

532

JOURNAL OF CLIMATEVOLUME 27

FIG. 2. The 5-yr average 2057–61 (left) September ice concentration and (right) JASON ice thickness from theindividual ensemble members with the (top) largest and (bottom) smallest 40-yr trend in ice EXT or ice VOL,
respectively. The ensemble-mean 15% ice concentration and 0.5-m ice thickness isopleths are contoured in black.

progresses. Mechanisms other than those directly associated with GHG forcing and the associated feedbacks
contribute importantly to the magnitude of 40-yr Arctic
sea ice loss trends in CCSM3.
Sea ice concentration and thickness trend maps arealso shown in Fig. 1 for the ensemble members with thelargest and smallest EXT and VOL trends. Althoughsubstantial ice concentration and thickness losses occur
in all ensemble members, they tend to be smaller andconcentrated around Greenland and the CanadianArchipelago in the ensemble members with the least sea ice
loss (Fig. 1, bottom maps) and larger and spread throughout the central and eastern Arctic in the ensemble members with the most sea ice loss (Fig. 1, top maps).
The amount and distribution of sea ice averaged overthe last five years of the simulation period (2057–61) areillustrated for the same ensemble members in Fig. 2.
Ensemble members with relatively large ice loss trends

exhibit mostly ice-free conditions in the central Arctic
by 2057–61 (Fig. 2, top). The ensemble members withrelatively small ice loss trends retain substantially more
extensive, more concentrated, and thicker sea ice inboth the central Arctic and the East Siberian Sea (Fig. 2,
bottom). These results illustrate that the concentrationand thickness of summer Arctic sea ice cover remaining
in the mid-twenty-first century is subject to considerable
uncertainty as a result of internal climate variability.
In summary, the magnitude of Arctic sea ice loss
projected to occur during 2020–59 is influenced not onlyby the buildup of GHGs but also by internal variability.
This internal variability can augment or diminish the
GHG-induced ice loss in any individual model realization. Intrinsic amplification mechanisms (e.g., positive
ice–albedo feedback) acting on randomly generated sea
ice anomalies do not dominate the spread in ice loss
among the 39 CCSM3 ensemble members.

15 JANUARY 2014

WETTSTEIN AND DESER

533

FIG. 3. Ensemble trend regressions between 2020–59 trends in September sea ice EXT and 2020–59 trends in (a) sea ice concentration,(b) sea ice thickness, (c) SLP, (d) U200, (e) Z500, and (f) Z1000. Regressions of (b) sea ice motion and (c) winds in the lowest (surface)atmospheric model layer are also shown along with reference vectors for scale. All ensemble trend regression calculations are based on
annual-mean values from October to September to reflect the sea ice year. The EXT trend values are multiplied by 21 so that larger values
reflect increased ice loss [see (a) and (b)]. All plotted regression coefficients are per standard deviation (0.83 3 106 km2) of the ensemble
spread in 2020–59 September Arctic sea ice EXT trends (i.e., the histogram from top left in Fig. 1). The 2020–59 time-mean and ensemblemean climatology of annual U200 is contoured in (d) at 5 m s21 contour intervals starting at 30 m s21 to indicate the location of themidlatitude jet at 200 hPa. Top and bottom panels show relationships poleward of 688 and 308N, respectively.

b. Ice loss, ice advection, and atmospheric circulation
Ensemble trend regressions are a compact and systematic way to explore relationships between internalvariability in the trends of different variables with thosein ice loss. Ensemble trend regressions onto EXT are
shown for a variety of annually averaged ice and atmospheric variables in Fig. 3. For example, Fig. 3c shows
the 40-yr SLP trends that are linearly associated witha one standard deviation change in the 40-yr trend of
September Arctic sea ice EXT. Ice concentration and

ice thickness ensemble trend regressions (Figs. 3a,b) arelargest off the coast of Eurasia and in the eastern Arctic,
a pattern that qualitatively matches the spatial distribution of concentration and thickness trend differences
between the ensemble members with the largest andsmallest ice loss trends (recall Fig. 1). Ice velocity ensemble trend regressions onto EXT indicate a coherent
pattern of sea ice advection that originates off the Siberian coast, crosses the central Arctic, and exits the Arctic
through the Fram Strait and Barents Sea (Fig. 3b). This‘‘transpolar’’ ice advection is coherent with near-surface

534

JOURNAL OF CLIMATEVOLUME 27

FIG. 4. As in Fig. 3, but for 2020–59 trends in annual (a) ice thickness and ice velocity, (b) SLP and near-surface wind, and (c) Z500 fields
regressed onto the standardized ensemble spread in 2020–59 JASON Arctic sea ice VOL trends instead of September sea ice EXT trends.
Reference vectors indicating a regression value of 1 cm s21 and 0.5 m s21 per standard deviation (2.7 3 103 km3) are shown for scale next to
the ice velocity and surface wind regression maps, respectively. Relationships poleward of 688N are shown in (a) and (b) and relationships
poleward of 308N are shown in (c).

atmospheric winds that are in turn associated witha SLP gradient between the western and the eastern
Arctic (Fig. 3c). The ensemble trend regression in SLP
is reminiscent of the dipole anomaly/Arctic dipole(AD) described by Wu et al. (2006) and Overland andWang (2010). The pattern of ice thickness ensemble
trend regression coefficients appears to be associatedwith the ice velocity field in the central Arctic, as evidenced for example by the two lobes in the ice thickness regression map that straddle the strongest central
Arctic ice velocities and also by the weak tongue of ice
thickness regression coefficients extending toward theFram Strait region of ice export.
Expanding our analysis domain to 308N, ensemble
trend regressions illustrate a coherent pattern of atmospheric circulation covariability with EXT from the levelof the subtropical jet (U200, Fig. 3d), through the midtroposphere (Z500, Fig. 3e), and down to the surface(Z1000, Fig. 3f). Extratropical atmospheric covariability
is concentrated into a geopotential height dipole patternover the Pacific sector with centers of action in the
subpolar Arctic and in the midlatitude North Pacific(Figs. 3e,f). The vertical structure of this pattern appears
to be largely equivalently barotropic (Figs. 3d–f). (Corresponding ensemble trend correlations with EXT are
shown in Fig. A1 of the appendix.)
A subset of the ensemble trend regression maps
from Fig. 3 are repeated in Fig. 4 using VOL trends as
the index variable. Ice advection in these VOL-based

regression maps is qualitatively similar to the analysis
based on EXT, but sea ice velocity regression coefficients are relatively more concentrated into a gyrelike circulation near the North Pole. In addition, sea ice
thickness regression coefficients are stronger throughout the central Arctic and extend more obviously intoa Fram Strait tongue (Fig. 4a). The dipole pattern in the
SLP ensemble trend regression onto VOL places moreemphasis on the eastern side of the Arctic, but themagnitudes of the SLP gradient and associated surface
wind regressions across the central Arctic (Fig. 4b) arecomparable to those based on EXT (Fig. 3c). The Z500
ensemble trend regression onto VOL (Fig. 4c) is slightly
stronger over the central North Pacific than the comparable regression onto EXT (Fig. 3e). Overall, the Z500
ensemble trend regression onto VOL is suggestive of
a Rossby wave train emanating from the midlatitude
central Pacific. All of the atmospheric ensemble trend
regression patterns associated with VOL are rotatedroughly 308 counterclockwise (eastward) relative tothose associated with EXT (cf. Figs. 4b,c with Figs. 3c,e).
In general, the Z500 ensemble trend regression onto
VOL in Fig. 4c provides a more compelling link to the
midlatitudes than the analogous regression onto EXT
in Fig. 3e in that the North Pacific lobe is more intense
and more coherently linked to a downstream Rossby
wave train. We will return to a comparison of the EXTand VOL patterns of atmospheric covariability in thediscussion.

15 JANUARY 2014

WETTSTEIN AND DESER

535

FIG. 5. (a) Seasonality of 2020–59 SLP ensemble trend regressions onto EXT (see text for details). Pattern correlation magnitudes are indicated on the left axis and pattern regression magnitudes normalized to the annual-mean
AD regression map (Fig. 3c) are indicated on the right axis. AMJJ regressions of (b) SLP and near-surface wind, (c)sea ice thickness and ice velocity, and (d) Z500 (color shading) are shown per standard deviation (0.83 3 106 km2) ofthe ensemble spread in 2020–59 September Arctic sea ice EXT trends. Corresponding AMJJ ensemble trend correlation coefficients between Z500 and EXT are also shown [see (d); contours] to enhance comparability with Figs. 6band A1e. Relationships poleward of 688N are shown in (b) and (c), whereas relationships poleward of 308N are shownin (d).

To assess the seasonal dependence of the atmosphericcirculation ensemble trend regressions on EXT (andvolume), we compute area-weighted pattern correlations and pattern regressions between the annual ensemble trend regression maps of Fig. 3c (Fig. 4c) andsimilar ensemble trend regression maps calculated using
monthly atmospheric trends. Pattern correlations andregressions are performed poleward of 708N for SLP andpoleward of 308N for Z500.

The annual-mean Arctic dipole–like SLP association
with EXT (Fig. 3c) is concentrated between April andJuly (AMJJ; Fig. 5a), and the corresponding AMJJ SLP
and surface wind ensemble trend regression patterns(Fig. 5b) are roughly twice as strong as in the annual
mean (Fig. 3c). The transpolar drift in ice velocity (Fig. 5c)is also more intense during AMJJ and more clearly associated with ice advection away from the Eurasian coast
than in the annual average (Fig. 3b). Finally, the AMJJ

536

JOURNAL OF CLIMATEVOLUME 27

FIG. 6. (a) Seasonality of 2020–59 Z500 ensemble trend regressions onto VOL (see text for details). Pattern correlation magnitudes are indicated on the left axis and pattern regression magnitudes normalized to the annual-mean
Rossby wave train regression map (Fig. 4c) are indicated on the right axis. (b) December–February (DJF) regressionsof Z500 (color shading) are shown per standard deviation (2.7 3 103 km3) of the ensemble spread in 2020–59 JASON
Arctic sea ice VOL trends. Corresponding DJF ensemble trend correlation coefficients between Z500 and VOL are also
shown (contours) to enhance comparability with Fig. 5d. Relationships poleward of 308N are shown in (b).

Z500 ensemble trend regression and ensemble trend correlation patterns (Fig. 5d) emphasize the central Arctic
lobe of a Pacific Rossby wave train compared to the annualmean regressions (Fig. 3e) and correlations (Fig. A1e). It
is noteworthy that the ensemble trend regression of
AMJJ Z500 on EXT (Fig. 5d) bears considerable similarity to the observed Z500 anomalies in the months preceding the previous record minimum Arctic sea ice extent
in 2007 [e.g., Fig. 3a in L’Heureux et al. (2008)].
The annual-mean association between VOL trends
and the Z500 Rossby wave train in the Pacific sector(Fig. 4c) is concentrated from December to February(DJF), with a secondary maximum in the pattern correlation during June. However, the pattern regression(i.e., the relationship between the annual and monthly
ensemble trend regression patterns in physical units) is
substantially weaker in June than in DJF (Fig. 6a). The
DJF Z500 ensemble trend regression pattern exhibits
a pronounced Pacific Rossby wave train (Fig. 6b) thatis more than twice as strong as in the annual average(Fig. 4c). The interpretation of this pattern as a Rossby
wave train emanating from the Pacific is supported by
the fact that the strongest center of action in Fig. 6boccurs over the North Pacific and weaker centers occur
downstream over northwest Canada and off the east
coast of North America.

It is worth noting that although the ensemble trend
regressions of Z500 onto VOL emphasize the Rossby
wave train pattern during the DJF season, weaker but
consistently positive pattern correlations and regressions
exist in all months of the year (Fig. 6a). Compared to
DJF, the Z500 Rossby wave train pattern associated withextent trends during AMJJ (Fig. 5d) exhibits a shorter
length scale (higher wavenumber), and its strongest centerof action is in the Arctic instead of the North Pacific.

c. Interannual atmospheric forcing of sea ice loss in
CCSM3, observations, and CCSM4
Prior subsections have described covarying trends ofsea ice and atmospheric circulation within the 39-member
CCSM3 ensemble. This subsection explores to what extent similar relationships may be found at interannual
time scales not only within the CCSM3 ensemble, butalso in observational products and within smaller CCSM4
ensembles.
Lead–lag bimonthly regressions of Z1000, Z500, ice
velocity, and ice thickness onto a 1-yr differenced VOL
time series are shown in Fig. 7. Each regression map isbased on 1560 time points resulting from the concatenation of time series during 2020–59 across the 39 ensemble members. The 1-yr differencing serves as aneffective high-pass filter to remove the strong trend

15 JANUARY 2014

WETTSTEIN AND DESER

537

FIG. 7. Lead–lag 2020–59 bimonthly regressions onto standardized 1-yr differenced values of the JASON ice volume. (top) The Z1000(color shading) and Z500 (contours; contour interval 4 m) regressions onto the JASON ice volume time series. (bottom) Ice velocity(vectors) and 1-yr differenced ice thickness (color shading) regressions onto the same JASON ice volume time series. Time advances as
one reads column by column from upper left to lower right. Regressions are for the 2-month averages indicated in the upper left of each
top row panel (e.g., August–September is denoted AS). Regressions leading and in the year prior to the JASON VOL index are denotedwith a negative sign (e.g., A-S2 indicates results are based on 2019–58 August–September averages). The JASON months that aresynchronous with the index are indicated with an asterisk. The JASON ice VOL time series and all spatial regressions are composed of1560 time points (40 yr 3 39 ensemble members). The contour interval for Z1000 (m) and ice thickness (cm) regressions is indicated below
the colorbar. Reference vectors indicating 1 cm s21 ice velocity regressions are indicated at the bottom right of each panel containing such
regressions. Vectors representing ice velocities below 0.2 cm s21 and the spurious ice velocity regressions at the sea ice edge are excluded
for clarity. Geopotential height regressions are shown poleward of 308N, whereas ice velocity and thickness regressions are shownpoleward of 688N.

538

JOURNAL OF CLIMATE

component from the raw VOL time series, and it is also
physically motivated, given that the observed time tendency in the amount of summer Arctic sea ice is influenced by the overlying atmospheric circulation andassociated with wind-driven ice advection (e.g., see papers cited in the introduction). The 1-yr differencing of
the ice thickness time series at each gridpoint and regressing these onto the 1-yr differenced VOL time series
provides an indication of the evolving spatial distribution of high-pass filtered ice thickness anomalies associated with high-pass filtered VOL changes. Unfilteredand high-pass filtered Z1000, Z500, and ice velocity regressions onto the 1-yr differenced VOL time series
yield similar results (not shown); the results based on
unfiltered data are displayed in Fig. 7 as a morestraightforward indicator of the patterns of atmosphericcovariabilty (Z1000 and Z500) and the response in ice
advection (ice velocity).
The spatial patterns and seasonal evolution of thehigh-pass filtered atmospheric circulation regression
maps are similar to those of the ensemble trend regression
maps (e.g., compare Fig. 7 with Figs. 5 and 6). In particular, the predominance of negative geopotential height
regression coefficients over the North Pacific with downstream centers of action resembling a Rossby wave train
in December through March evolves into an Arcticconcentrated mostly annular structure by June–July(Fig. 7). A weak equivalent barotropic North Pacific low
pressure center exists already in October–November ofthe year preceding the summer sea ice VOL index.
A weak Arctic dipole–like pattern in central Arctic
Z1000 regression coefficients is also already present in
October–November, and this pattern persists throughFebruary–March and then strengthens substantially in
April–May. Positive geopotential height regression coefficients fill the Arctic basin in June–July, but a similar
east–west SLP gradient persists across the Arctic because the strongest positive SLP anomalies are positioned over the western Arctic. After July, central Arctic
geopotential height regression coefficients weaken substantially. Note that the evolution in atmospheric circulation associated with Fig. 7 is strictly only related to sea
ice covariability and that a separate analysis of the atmospheric variability would be necessary to link theevolution of the various patterns dynamically.
Ice export across the central Arctic through the Fram
Strait and into the Barents Sea occurs throughout theyear preceding the summer sea ice VOL index, but the
distribution of the central Arctic ice velocity pattern
evolves through the year (Fig. 7). The evolving ice velocity pattern is consistently and coherently linked to the
distribution and intensity of near-surface winds (notshown but they can be inferred from the Z1000 gradient).

VOLUME 27

Ice velocity seems to lead associated anomalies in ice
thickness by a month or two (cf. the divergence of the ice
velocity field with the development of ice thickness
anomalies in the subsequent bimonthly regression map),suggesting that ice advection influences the overall
evolution of ice thickness anomalies in the central
Arctic. For example, the divergence of the ice velocity
field especially in the western Arctic during April–May
can be associated with thinner ice in the same area
during June–July. Ice export from the central Arctic
ceases by the end of September, consistent with theslackening of surface winds and weaker Z1000 gradients.
The pattern of ice thickness anomalies during October–
November at the end of the bimonthly lagged regressions (lower right panel of Fig. 7) is qualitatively similar
to the pattern resulting from ensemble trend regressionsof ice thickness onto VOL in Fig. 4a.
The results shown in Fig. 7 can be summarized in a few
key points. First, the patterns and seasonality of atmospheric circulation covariability with Arctic sea ice loss
are generally similar between high-pass filtered andtrend-based analyses. Second, ice advection out of thecentral Arctic is a consistent feature in the bimonthly
lead–lag analysis and is itself directly related to theoverlying Arctic atmospheric circulation. Third, ice divergence (convergence) can often be related to subsequent sea ice thickness decreases (increases). Finally,
the evolution of sea ice–atmospheric circulation covariability is strongest in the year preceding the summer
minimum in Arctic sea ice and becomes relatively weakby the last (contemporary) months of the analysis.
Qualitative similarities in the seasonality and spatial
distribution between VOL-based regressions using highpass filtered data and 40-yr trends suggest that a similar
mechanism of wind-driven ice advection is operatingover a broad range of time scales. Because of this, we
explore the relationships between ice motion and SLP
anomalies in the short observational record to validate
the model results. Figure 8a shows the observed 1979–2010 September sea ice extent time series in both anunfiltered form (black curve) and after a 1-yr difference
filter has been applied (red curve; note that extent isused rather than volume because observational estimates of the latter are less robust than those of theformer). AMJJ regressions of near-surface winds and
SLP from the ERA-Interim reanalysis and of observed
ice velocity onto the observed 1-yr differenced September
Arctic sea ice extent time series are shown in Figs. 8b and8c, respectively. The spatial pattern and seasonality of theobserved SLP and wind regressions are broadly similar to
those identified by Screen et al. (2011). The large-scalepatterns in these observed AMJJ regression maps also
bear some resemblance to AMJJ regressions of the same

15 JANUARY 2014

WETTSTEIN AND DESER

FIG. 8. (a) The unfiltered 1979–2010 September Arctic sea ice extent time series obtained from the NSIDC (black
curve) is shown along with the 1-yr differenced time series of the same (red curve). (b) The 31-yr (1980–2010)
regression of AMJJ SLP and 1000-hPa winds from the ERA-Interim (ERA-I) reanalysis is shown along with (c) the
27-yr (1980–2006) regression of AMJJ ice velocity from NASA’s Polar Pathfinder onto a standardized inverse timeseries of the red curve in (a). (d),(e) A similar CCSM3 regression onto 1-yr differenced September EXT using 1560
time points from 2020–59 and concatenated across the 39 CCSM3 ensemble members is shown. Ice thickness regression coefficients are included in the CCSM3 analysis shown in (e).

539

540

JOURNAL OF CLIMATEVOLUME 27

FIG. 9. Regressions of 1-yr differenced AMJJ sea ice thickness (color shading) and unfiltered ice velocity (vectors) onto standardized
1-yr differenced JASON ice VOL (a) in the CCSM3 ensemble and (b),(c) in two of the four CCSM4 RCP ensembles. The CCSM4
ensembles range from relatively weak equivalent GHG forcing (RCP2.6) in (b) to relatively strong equivalent greenhouse gas forcing(RCP8.5) in (c). Each map is based on a concatenation across 40 yr of data and the number of ensemble members indicated in parenthesis.
The climatological-, ensemble-, and annual-mean sea ice thickness from the same 40 yr is contoured in green at 1-m intervals up to 5 m.
The 1- and 5-m contours are labeled for clarity.

fields onto the concatenated 1-yr differenced EXT timeseries from CCSM3 (Figs. 8d,e), though differences are
sometimes large locally. The AMJJ association withhigh-pass filtered CCSM3 EXT time series (Figs. 8d,e)can also be compared with the April–May and June–
July regression maps onto high-pass filtered CCSM3
VOL time series (Fig. 7, lower left). This comparison
indicates that the AMJJ ice thickness regression onto
high-pass filtered EXT (Fig. 8e) is associated with asimilar seasonal evolution of ice advection to thatshown for VOL in Fig. 7. For example, the negative ice
thickness anomalies in the Beaufort Sea, the positive
anomalies downstream of Wrangel Island, and thepositive anomalies in the Fram Strait (Fig. 8e) can all
be associated with the evolution of wind-driven sea ice
convergence and divergence in the months prior to
AMJJ. Although there is an Arctic dipole–like AMJJSLP gradient associated with Arctic sea ice extent
variability in both ERA-Interim (Fig. 8b) and CCSM3(Fig. 8d), the CCSM3 SLP pattern strongly emphasizes
the high over the western Arctic relative to the low
over the eastern Arctic. The large-scale pattern of ice
motion in CCSM3 (Fig. 8e) is also rotated roughly 308
counterclockwise (eastward) relative to the large-scalepattern of observed NASA Pathfinder ice motion regressed onto observed sea ice extent (Fig. 8c). A transpolar drift in ice velocity culminating in sea ice export
through the Fram Strait and into the Barents Sea is,however, a common feature associated with both observed and CCSM3-simulated summer high-pass filtered

Arctic sea ice variability, and this extends to variability of
40-yr trends in CCSM3.
Regressions of AMJJ ice velocity and 1-yr differenced
AMJJ ice thickness onto 1-yr differenced VOL timeseries in CCSM3 and CCSM4 are directly compared inFig. 9. CCSM3 regressions of AMJJ ice velocity and 1-yr
differenced ice thickness on 1-yr differenced VOL timeseries (Fig. 9a) are as expected based on the April–May
and June–July bimonthly regressions (Fig. 7, lower left).
CCSM4 regressions of the same AMJJ variables in ensembles from both the relatively weak equivalent GHG
forcing (RCP2.6; Fig. 9b) and the relatively strong
forcing (RCP8.5; Fig. 9c) and also in the intermediate
RCP4.5 and RCP6.0 ensembles (not shown) are all associated with a transpolar drift in ice velocity andthickness anomalies that are consistent with an evolution of ice advection away from the coast, across thecentral Arctic, and out through the Fram Strait and into
the Barents Sea.
CCSM4 simulations exhibit a climatological-mean ice
thickness distribution that better matches available observations (green contours in Fig. 9; see also Jahn et al.
2012) with thicker ice in the western Arctic comparedwith the T42 resolution CCSM3 ensemble. The T42
resolution CCSM3 ensemble simulates a bimodal ice
thickness distribution with maxima surrounding both
the Canadian Archipelago and the coast of the East
Siberian Sea. Although AMJJ ice advection and (implied) atmospheric surface pressure regressions are differently distributed in CCSM4 and observations than in15 JANUARY 2014

WETTSTEIN AND DESER

CCSM3, volume loss in each dataset is associated with
ice advection away from the coast in the regions ofthickest climatological-mean sea ice, a transpolar drift in
ice advection through the central Arctic, and ice export
through the Fram Strait and into the Barents Sea. Thus,
the rotated Arctic dipole–like pattern in CCSM3 compared to observations (Fig. 8) and compared to CCSM4(Fig. 9) may result from the erroneous climatologicalmean distribution of Arctic sea ice in the T42 resolution
CCSM3, rather than any particular deficiency in thesimulated physical mechanism relating internal variability in Arctic sea ice loss to atmospheric circulation.
Equivalently, the better agreement between CCSM4
and the observed distribution of ice advection and (implied) atmospheric surface pressure and surface windfield regressions onto high-pass filtered sea ice variability may result from the better simulation of theclimatological-mean Arctic sea ice distribution in CCSM4
than in the T42 CCSM3.

d. Atmospheric relationships with internal sea ice
trend variability: Intrinsic patterns and global
linkages
Previous subsections have shown that internal variations in Arctic ice loss trends are associated both with an
Arctic dipole–like pattern over the central Arctic andalso with atmospheric pressure anomalies over the
North Pacific. It is interesting to consider whether these
constitute intrinsic patterns of variability in their own
right, regardless of their association with sea ice. EOF
analysis is used to characterize the leading patterns ofannual-mean and seasonal-mean Z500 and Z1000 trend
variability in the CCSM3 ensemble (Fig. 10, color shading). The first EOFs of both annual and December–
February North Pacific Z500 trend variability (Fig. 10,top) are associated with Rossby wave train patterns thatbear considerable similarity to, but are stronger than, thecomparable Z500 ensemble trend regressions onto VOL.
Leading EOFs of both annual and April–July Z1000 trend
variability poleward of 708N are associated not only withan AD-like pattern over the central Arctic, but also covarying pressure anomalies over the North Pacific(Fig. 10, bottom). The Z1000 EOFs in Fig. 10 bear considerable similarity to the comparable Z1000 ensemble
trend regressions onto EXT. Area-weighted patterncorrelations poleward of 308N between the trend EOFs
(color shading) and the ensemble trend regressions onto
VOL or EXT (contours in the top and bottom panels,
respectively) are in all cases high, ranging between 0.81
and 0.92. It is particularly noteworthy that the 39 PC2
values representing April–July Z1000 trend variability(Fig. 10, bottom right) share half of the variance (r 50.71) with the 39 different values for the 2020–59 EXT

541

trend in the CCSM3 ensemble. Note that the third EOF of
Z1000 annual trends and the second EOF of Z1000 AMJJ
trends are displayed, but that each EOF in Fig. 10 is statistically distinct from all other EOFs when the first-ordereigenvalue uncertainty estimate described in North et al.
(1982) is applied with 39 degrees of freedom. The first
EOF of both annual and AMJJ Z1000 trends is an Arctic
Oscillation–like pattern and the second EOF of annual
Z1000 trends is a dipole over the Arctic that is mostly associated with North Atlantic variability (not shown). EOF
patterns similar to those in Fig. 10 are obtained using Z500
and Z1000 trends over different spatial domains and within
other seasons (not shown). Comparable EOFs are also
obtained using a 1000-yr CCSM3 control integration (notshown) and have been obtained in EOF-based analysis ofinterannual variability in reanalysis (e.g., Overland andWang 2005; Wu et al. 2006; Wang et al. 2009).
It is worth noting that the trend-based EOF analysis isdesigned to pick out the leading patterns of annualmean or seasonal-mean trend variability across the 39
ensemble members independent of any Arctic sea ice
relationships. The similarity between these ensemble
trend EOFs and the ensemble trend regressions used to
identify atmosphere–sea ice covariability suggests thatinternal variability in the trends of Arctic sea ice loss
is associated with naturally occurring patterns of lowfrequency atmospheric circulation variability.
The global signature of atmospheric circulation trends
associated with internal variability in VOL trends
is depicted in Figs. 11 and 12 (similar results are obtained using EXT trends, not shown). Annual-meangeopotential heights averaged over the Pacific sector(1208E–908W) exhibit ensemble trend correlations andregressions with VOL that extend coherently throughout
the depth of the troposphere, with positive values poleward of 558N and negative values in the midlatitudes(Fig. 11). A weak extension of geopotential height covariability into the tropics is also evident, suggestinga linkage between internal variability in trends of Arctic
sea ice and tropical Pacific atmospheric circulation.
Global ensemble trend regression and correlation
maps of annual-mean Z500, SLP, and Tref onto VOL(left panels, Fig. 12) confirm the extension of ice-related
atmospheric trends into the tropics and into the Southern Hemisphere. In particular, internal variability in
VOL trends is associated with positive Z500 trends
throughout the tropics and a SLP trend dipole between
the tropical Pacific and Atlantic. A baroclinic structure
is evident in the atmospheric pressure anomalies overthe central tropical Pacific (cf. Figs. 12a,c). In addition,
positive Z500 and SLP trends are found over the Pacificsector of the Southern Ocean. These tropical and extratropical features associated with Arctic ice VOL loss

542

JOURNAL OF CLIMATEVOLUME 27

FIG. 10. Leading EOFs (color shading) of 2020–59 trends in (left) annual or (right) seasonal (top) Z500 or (bottom)
Z1000 are calculated from the 39-member CCSM3 ensemble and displayed as regressions onto standardized values of
the associated PCs. The Z1000 and Z500 EOFs are calculated for the AMJJ and DJF seasons during which the EXT
associations with Z1000 or the VOL associations with Z500 are strongest (Figs. 5 and 6, respectively). The same annual
or seasonal Z500 and Z1000 trends are also regressed (contours) onto the 39 CCSM3 VOL or EXT trend values as
indicated. The Z500 trend PCs are calculated poleward of 308N in an extended Pacific sector (1208E–908W) indicatedby a green line. The Z1000 trend PCs are calculated poleward of 708N as indicated by the green circle. The contour
interval for both maps in each panel is indicated in the lower left.

are accompanied by increased convective precipitationand latent heating over most of the tropical Pacific (notshown). The global structure of internal variability inatmospheric circulation trends associated with Arctic ice
VOL can also be obtained from the leading PC of
tropical Pacific Tref trends (right panels, Fig. 12). Thispattern is reminiscent of a ‘‘global hyper mode’’(Dommenget and Latif 2008) that may result fromthermodynamic air–sea interaction (see also Clement et al.
2011).

The relevance of the tropical Pacific in internal Arctic
sea ice loss variability is demonstrated more directly in
Fig. 13, which provides maps comparable to those in
Figs. 3a–c, but for annual ice and atmosphere trend regressions onto the first PC of tropical Pacific Tref trends
instead of onto EXT trends. The regression maps show
an Arctic dipole–like SLP pattern and associated nearsurface winds (Fig. 13c) that drive a transpolar drift in
ice velocity and ice advection out of the Arctic throughthe Fram Strait and into the Barents Sea (Fig. 13b). The

15 JANUARY 2014

WETTSTEIN AND DESER

FIG. 11. Ensemble trend correlations (color shading) andensemble trend regressions (contours) between annual (October–
September) 2020–59 trends in the Pacific (1208E–908W) zonal average geopotential height field and standardized 2020–59 JASON
Arctic sea ice VOL trends. Thicker black contours represent a
Pacific correlation/regression value of zero. Dashed contours indicate negative regression values. The regression contour interval
is 1 m per standard deviation (2.7 3 103 km3) of the JASON VOL
trends.

first PC of tropical Pacific Tref trends is also associatedwith negative sea ice concentration (Fig. 13a) andthickness (Fig. 13b) poleward of Alaska and Siberia.

4. Summary
This study’s results can be summarized into the following key findings:
Internal variability is an important factor influencing
the magnitude and spatial pattern of twenty-firstcentury summer Arctic sea ice loss in the 39-member
CCSM3 ensemble forced by the A1B GHG emissions
scenario. Specifically, internal variability leads toa roughly threefold range in trends of ice EXT and
VOL loss during 2020–59.
d This spread in the amount of ice loss is related to largescale patterns of atmospheric circulation and accompanying ice motion fields.
- In particular, an Arctic dipole–like pattern in SLP
trends that is strongest in spring and early summer(April–July) is associated with Arctic sea ice advection across the central Arctic and export through theFram Strait and into the Barents Sea.
- Internal variability in sea ice loss trends is also
associated with an atmospheric Rossby wave train
pattern in the free troposphere, with centers ofaction over the North Pacific and downstream overd543

northwestern North America and the western Atlantic. This Rossby wave train is emphasized during
winter (December–February) but is present throughout the year.
d Each of these atmospheric pressure patterns is associated with a leading pattern of atmospheric trend
variability within the ensemble, as determined throughindependent EOF/PC analysis.
d The atmospheric circulation patterns associated withinternal variability in Arctic sea ice loss exhibit weak
but significant teleconnections to the tropical andSouthern Hemisphere Pacific in annual-mean data.
Similar Pacific teleconnection patterns, with extensions over the Arctic, are associated with the leading
PC of Tref trends over the tropical Pacific.
d Associations between sea ice and atmospheric circulation anomalies on interannual time scales arebroadly similar to those that characterize the internalvariability of 40-yr trends.
- In particular, the spatial pattern and seasonal dependence of the atmospheric circulation anomalies
that covary with high-pass filtered sea ice variability as well as the accompanying changes in transArctic sea ice advection resemble those based oninternal trend variability.
- In CCSM3, high-pass filtered variability in Arctic
sea ice loss is associated with ice advection throughthe Fram Strait and into the Barents Sea throughout the year preceding the summer Arctic sea ice
minimum.
- Interannual relationships between sea ice and atmospheric circulation anomalies simulated by CCSM3
are generally realistic except for a 308 rotation of thelarge-scale wind-driven ice motion across the central Arctic. This shortcoming is likely related toa bias in the CCSM3’s climatological-mean sea ice
thickness distribution, an aspect that is improved in
CCSM4.

5. Discussion
The results of this study describe an important role forinternal variability in twenty-first-century projections ofArctic sea ice loss and illustrate a consistent set ofphysically related patterns in the distribution of ice loss,
ice motion, and atmospheric circulation variability. This
discussion will elaborate on and provide an interpretation of some of the results.
Summer Arctic sea ice extent (Fig. 14a) and volume(Fig. 14b) projections from the smaller (e.g., 6–7 member) CCSM4 climate change ensembles (the same thatwere used to generate Figs. 9b,c) demonstrate that internal variability remains an important component of544

JOURNAL OF CLIMATEVOLUME 27

FIG. 12. Ensemble trend correlations (color shading) and ensemble trend regressions (contours) with (left)
JASON VOL and with (right) the first PC of annual (October–September) tropical Pacific near-surface reference temperature (Tref, PC1) are shown. Correlations with annual (October–September) 2020–59 trends in (a),(b) Z500, (c),(d) SLP, and (e),(f) Tref indicate the global-scale atmospheric teleconnections with (left) Arctic iceloss and (right) tropical Pacific temperature. The tropical Pacific (208S–208N, 1208E–908W) domain over which
the Tref PCs were calculated is indicated by a green box in (f). Contour intervals for the regressions are indicatedin the lower left of each panel.

uncertainty in the newest version of the model (see also
Kay et al. 2011) as it is in the CCSM3 ensemble weanalyzed. In particular, Fig. 14 suggests that internalvariability will continue to be an important factor in
future Arctic sea ice loss projections for at least the next
several decades and that this conclusion is robust, notan artifact produced only within the CCSM3 ensemble.
For instance, one CCSM4 ensemble member from theweakest equivalent GHG forcing scenario (RCP 2.6) is
associated with the least extensive September sea ice aslate as 2051, while a member from the second strongest
equivalent GHG forcing scenario (RCP 6.0; similar to
the SRES A1B) is associated with the most extensive
September sea ice several times between 2052 and 2061.
These two examples demonstrate the lingering influenceof internal variability in CCSM4. A qualitative assessment of uncertainty in Arctic sea ice loss projections using other coupled models in the CMIP3
archive suggests that the three sources of uncertainty(magnitude of GHG forcing, model sensitivity, andinternal variability) are all roughly comparable in the
middle of the twenty-first century (not shown). This
study describes an important role for internally generated variability in the magnitude of twenty-firstcentury Arctic sea ice loss; a generally similar result
was obtained in a comparable but much smaller(5 member) coupled model ensemble (Sorteberg et al.
2005). Additional ensemble members with other fully
coupled models are required for a more robust
evaluation.

15 JANUARY 2014

WETTSTEIN AND DESER

545

FIG. 13. As in the top panels of Fig. 3, but the trend regressions are onto the first PC of tropical Pacific near-surface temperature trends(Tref PC1) instead of onto Arctic EXT trends. See Fig. 12 and the text for more detail on the tropical Pacific Tref PC calculation.

The qualitative similarity of the results based on theinternal variability of trends and that of the interannual
variability suggests that similar atmospherically driven
ice loss mechanisms operate over a broad range of timescales. Based on this similarity, we interpret the trendbased results as a natural consequence of large-scaleatmospheric circulation variability that exists andapparently drives changes in Arctic sea ice cover overa broad range of time scales.
As alluded to in the introduction, distinguishingwhether central Arctic atmospheric covariability withsea ice loss is more Arctic dipole (AD)–like or more
Arctic Oscillation (AO)–like is an active area of research, but no clear consensus exists. Although the

FIG. 14. Simulated twenty-first-century CCSM4 (a) September sea ice extent and (b) JASON sea ice volume areplotted as a function of CMIP5 RCP emissions scenarios indicated. The ensemble mean (thick solid lines) is shown
along with the ensemble range (color shading) as represented by the minimum and maximum September sea ice
extent or JASON sea ice volume value in each year for each RCP. The number of ensemble members is indicated inparenthesis next to the RCP scenario legend in the bottom left. The CCSM4 ensembles range from relatively weakequivalent GHG forcing (RCP2.6) to relatively strong equivalent GHG forcing (RCP8.5). More detail on the
CCSM4 ensemble and the RCP scenarios is contained in the text.

546

JOURNAL OF CLIMATE

analyses in this paper have emphasized an Arctic
dipole–like pattern of atmospheric covariability, results
have also documented generally stronger relationships
with ice advection and corresponding surface winds thanwith any static pattern of atmospheric covariability (e.g.,ensemble trend correlations with EXT are larger forsurface winds than for SLP; see appendix, Fig. A1c). A
similar conclusion has been distilled from the available
observations (Smedsrud et al. 2011). In particular, ouranalysis of interannual time-scale variability (e.g., Fig. 7)
and of intermediate time scales and composites of individual ice loss events (not shown) indicates that thepatterns of surface pressure anomalies associated withice loss evolve seasonally and can also change substantially from event to event, but that a transpolar driftin ice advection and Fram Strait/Barents Sea ice export
are consistently involved. Note, for example, that thecentral Arctic Z1000 regression pattern during June–Julyin Fig. 7 is associated with a negative AO-like pattern,but one whose polar center of action is concentrated
over the Canadian Archipelago in the western Arcticthat in turn affects surface winds throughout the Arctic
basin. In general, the collection of results in this paper
suggests a more flexible presumption about atmosphericcovariability with ice loss is warranted. Insofar as theAD-like pattern is associated with atmospheric variability emanating from the tropical Pacific in a Rossby wave
train (as it is in the CCSM3 ensemble), different flavors ofboth of the leading two patterns of extratropical surface pressure variability (the northern annular mode/AO
and the PNA patterns) may in turn be associated withanomalous ice advection, ice export, and internal variability in sea ice loss. A pattern of SLP variability similar
to the one we associate with internal variability in Arctic
sea ice advection and export was linked to multidecadal
climate variability in the Greenland Sea by Delworth
et al. (1997).
Internal variability in CCSM3 Arctic sea ice loss isrelated mostly to internal variability in the North Pacificatmospheric circulation and has relatively little association with atmospheric circulation variability over the
North Atlantic. The ensemble spread in atmosphericcirculation trends over the North Pacific is distinct fromthe spatial pattern of the ensemble-mean (forced) trend
over the North Pacific (not shown), so the covariability
between Arctic sea ice loss and North Pacific atmospheric circulation has a clear association with naturalvariability. The similarity between Pacific atmosphericcirculation covariability patterns with both Arctic ice
VOL trends and the leading PC of tropical Pacific Tref
trend variability suggests that the tropical Pacific coherently influences subtropical, midlatitude, and evenpolar circulation variability in both hemispheres. The

VOLUME 27

association between tropical Pacific variability andArctic atmospheric circulation is consistent with ice removal from the central Arctic via ice export through theFram Strait and into the Barents Sea.
Arctic sea ice loss in the CCSM3 ensemble is associated with the strength and longitudinal extent of theNorth Pacific subtropical jet (Fig. 3d) and with a Rossby
wave train emanating from the subtropical Pacific (Figs.
4c and 6b). Both of these signatures have been associated with the PNA and a leading pattern of observed
storm track–jet covariability in the North Pacific(Wettstein and Wallace 2010). The strength and longitudinal extent of the Pacific subtropical jet has also been
related to variability in the intensity of tropical diabatic
heating and the PNA (Li and Wettstein 2012). The
mostly equivalent barotropic midlatitude structure, thesimilarity of Pacific teleconnections between Arctic sea
ice loss and tropical Pacific Tref variability, and the linkto our understanding of midlatitude atmospheric dynamics all suggest that the internal variability we concentrate on in this paper is associated with an Arctic/
downstream response to tropical Pacific variability
rather than vice versa. Our results and interpretation are
also broadly consistent with the influence of the tropical
Pacific and a PNA-like pattern on Arctic surface airtemperature described by Lee et al. (2011).
A few of the CCSM3-based interpretations discussed
in the previous three paragraphs are worthy of special
attention.
(i) A transpolar drift in ice advection and Fram Strait–
Barents Sea ice export are consistently associatedwith internal variability in Arctic sea ice loss, even
if the precise SLP patterns vary by event and timeof year.
(ii) Internal variability in Arctic sea ice loss is related
to internal variability in the large-scale atmospheric circulation not only over the Arctic, butalso over the North Pacific (and not over the North
Atlantic).
(iii) Internal variability of the Arctic and North Pacificatmospheric circulation is influenced by the tropical Pacific via Rossby wave dynamics. Thus, tropical Pacific variability may play a role in anomalous
Arctic sea ice advection and export.
Both September ice EXT and July–November(JASON) ice VOL have been used in this study as indices of summer Arctic sea ice loss. Although the 40-yr
trends of EXT and VOL are highly correlated acrossthe 39-member ensemble (r 5 0.80), roughly a third of
the trend variance is unshared. Given that most of thetrend variance is common between the two indices ofsummer Arctic sea ice loss, qualitatively similar patterns

15 JANUARY 2014

547

WETTSTEIN AND DESER

and seasonality are obtained if an investigation of theseasonality of the Arctic dipole–like SLP association(e.g., Fig. 5) is based on VOL rather than EXT or if aninvestigation of the seasonality of the Rossby wave train
association (e.g., Fig. 6) is based on EXT rather thanVOL. That said, the particular pairing of SLP variability
with EXT and Z500 variability with VOL provides
stronger, more consistent, and more physically intuitive
patterns. Our interpretation of these results is that EXT
is somewhat more sensitive to the AMJJ wind-driven
ice advection, whereas VOL is responsive to a more
integrated annual average of sea ice advection and export (e.g., Figs. 6 and 7). A related interpretation can bemade for the roughly 308 counterclockwise pattern rotation of annually averaged atmospheric covariability
with VOL relative to EXT (cf. Figs. 3c,e with Figs. 4b,c).
Because VOL has a stronger relationship to annual sea
ice advection and to the Pacific Rossby wave train, andalso because the Pacific subtropical jet is stronger andzonally extended in winter (e.g., DJF) relative to latespring and early summer (e.g., AMJJ), the entire pattern of atmospheric covariability over the extratropical Pacific is shifted roughly 308 counterclockwise(downstream) in annual VOL regressions relative toannual EXT regressions. In general, the mechanism of
ice advection and export we emphasize and associatewith internal Arctic sea ice loss variability likely has
a stronger and more coherent association with sea ice
volume than with sea ice extent.
The large ensemble of twenty-first-century simulations by CCSM3 is currently a unique resource for exploring internal variability and its mechanisms, but there
are some noteworthy caveats. As mentioned previously,
the T42 resolution CCSM3 exhibits too much and improperly distributed Arctic sea ice relative to nature(DeWeaver and Bitz 2006). This bias could influence thesimulated sensitivity to atmospheric forcing. In addition,
like many CMIP3 models, CCSM3 generates El Ni~no–Southern Oscillation (ENSO) events that are too regular and too frequent, as well as a Pacific decadal oscillation (PDO) that lacks a strong connection between thetropics and the North Pacific. The extratropical atmospheric teleconnection patterns associated with both
ENSO and PDO at their native simulated frequencies
are nevertheless reasonable (Deser et al. 2006; Stoneret al. 2009).
By focusing on the sea ice response to atmosphericforcing, we largely neglect any feedbacks of the internalvariability in ice loss trends upon the atmosphere. Previous studies have shown that the atmospheric circulation response to Arctic sea ice anomalies is generally
much weaker than the atmospheric circulation patterns
that force sea ice changes (see Magnusdottir et al. 2004;

Deser et al. 2004; Kvamstø et al. 2004; Bhatt et al. 2008;
Seierstad and Bader 2009; Deser et al. 2010; Serreze
and Barry 2011; Bl€uthgen et al. 2012 and compare toreferences in the introduction). Preliminary results(not shown) indicate that the intensity of the SLP lowover the eastern Arctic may be related to the atmospheric response to Arctic sea ice loss in the CCSM3
ensemble, broadly consistent with the findings ofBl€uthgen et al. (2012). Isolating the response andfeedbacks to Arctic sea ice loss in the atmosphere,
ocean, and sea ice in observations and in the fully
coupled CCSM3 and CCSM4 ensembles is the subjectof our ongoing research.
The results in this study are based on one model(CCSM3), with additional supporting evidence fromCCSM4. Compared to CCSM3, CCSM4 shows an improved simulation of the distribution of Arctic sea ice(Jahn et al. 2012) and of the frequency spectrum of
ENSO (Deser et al. 2012c), highlighting new opportunities for mechanistic studies using large ensembles asmodels become more realistic. Large ensembles of climate change simulations at a variety of resolutions withother models are needed to test the robustness of ourfindings based on CCSM, in particular both the central
Arctic relationships and proposed associations with themidlatitude and tropical Pacific.
Acknowledgments. We thank Marika Holland, AlexJahn, Jen Kay, Adam Phillips, and Dennis Shea,
who have each contributed helpful suggestions on various aspects of this work. We would also like to thankJohn Walsh and two anonymous reviewers for theirthoughtful comments. We made extensive use of the
NCAR Command Language (NCL) and would like toacknowledge the National Science Foundation for theirsponsorship of this effort and especially Dennis Shea forhis support and guidance. J. W. was supported by a grantfrom the National Science Foundation Arctic Sciences
Program.

APPENDIX
Annual Ensemble Spread Correlations with Extent
This paper has analyzed various features of internalvariability in Arctic sea ice loss with a focus on atmospheric covariability. The annual ensemble trend regressions with EXT (Fig. 3) are probably the mostgeneral and comparable to other studies, but regression
coefficients demonstrate the relationships only in physical units, whereas the amount of shared variance is also
a useful quantity. Annual ensemble trend correlations

548

JOURNAL OF CLIMATEVOLUME 27

FIG. A1. As in Fig. 3, but ensemble trend correlations with EXT are shown instead of regressions.

with extent are shown in Fig. A1 for the same sea ice andatmospheric variables as in Fig. 3.
Ice concentration ensemble trend correlations with
EXT are strongest in the Eurasian sector, diminish inmagnitude across the central Arctic, and extend intoa tongue within the Fram Strait (Fig. A1a). This distribution is roughly similar for ice thickness (Fig. A1b), butcorrelation magnitudes are lower because the index
variable is EXT rather than VOL. The Arctic dipole–like ensemble trend correlation of SLP with EXT is
weak, but stronger correlations exist with the SLP gradient (not shown) and with near-surface winds (Fig.
A1c). A transpolar drift in ice velocity correlations is
consistent with wind-driven ice export arguments elsewhere in this paper. Large-scale atmospheric covariability is also demonstrated by significant ensemble
trend correlations outside the Arctic (Figs. A1d–f).
Geopotential height correlations can be compared withthose in Figs. 5, 6, 11, and 12.

REFERENCES
Belchansky, G. I., D. C. Douglas, V. A. Eremeev, and N. G.
Platonov, 2005: Variations in the Arctic’s multiyear sea ice cover:A neural network analysis of SMMR-SSM/I data, 1979–2004.
Geophys. Res. Lett., 32, L09605, doi:10.1029/2005GL022395.
Bhatt, U. S., M. A. Alexander, C. Deser, J. E. Walsh, J. S. Miller,
M. S. Timlin, J. Scott, and R. A. Tomas, 2008: The atmospheric
response to realistic reduced summer Arctic sea ice anomalies.
Arctic Sea Ice Decline: Observations, Projections, Mechanisms, and Implications, Geophys. Monogr., Vol. 180, Amer.
Geophys. Union, 91–110.
Bitz, C. M., 2008: Some aspects of uncertainty in predicting sea ice
thinning. Arctic Sea Ice Decline: Observations, Projections,
Mechanisms, and Implications, Geophys. Monogr., Vol. 180,
Amer. Geophys. Union, 63–76.
——, J. K. Ridley, M. Holland, and H. Cattle, 2012: Global climate
models and 20th and 21st century Arctic climate change.
Arctic Climate Change: The ACSYS Decade and Beyond,
P. Lemke and H.-W. Jacobi, Eds., Springer, 405–436.
Bl€uthgen, J., R. Gerdes, and M. Werner, 2012: Atmospheric response to the extreme Arctic sea ice conditions in 2007. Geophys. Res. Lett., 39, L02707, doi:10.1029/2011GL050486.

15 JANUARY 2014

WETTSTEIN AND DESERBo
e, J., A. Hall, and X. Qu, 2010: Sources of spread in simulationsof Arctic sea ice loss over the twenty-first century. Climatic
Change, 99, 637–645.
Branstator, G., and H. Teng, 2010: Two limits of initial-value decadal predictability in a CGCM. J. Climate, 23, 6292–6311.
Clement, A., P. DiNezio, and C. Deser, 2011: Rethinking theocean’s role in the Southern Oscillation. J. Climate, 24, 4056–
4072.
Collins, W. D., and Coauthors, 2006: The Community Climate
System Model version 3 (CCSM3). J. Climate, 19, 2122–2143.
Comiso, J. C., C. L. Parkinson, R. Gersten, and L. Stock, 2008:
Accelerated decline in the Arctic sea ice cover. Geophys. Res.
Lett., 35, L01703, doi:10.1029/2007GL031972.
Delworth, T. L., S. Manabe, and R. J. Stouffer, 1997: Multidecadal
climate variability in the Greenland Sea and surrounding regions: A coupled model simulation. Geophys. Res. Lett., 24 (3),257–260.
Deser, C., and H. Teng, 2008: Recent trends in Arctic sea ice andthe evolving role of atmospheric circulation forcing, 1979–
2007. Arctic Sea Ice Decline: Observations, Projections,
Mechanisms, and Implications, Geophys. Monogr., Vol. 180,
Amer. Geophys. Union, 7–26.
——, G. Magnusdottir, R. Saravanan, and A. S. Phillips, 2004: The
effects of North Atlantic SST and sea ice anomalies on thewinter circulation in CCM3. Part II: Direct and indirectcomponents of the response. J. Climate, 17, 877–889.
——, A. Capotondi, R. Saravanan, and A. S. Phillips, 2006: Tropical Pacific and Atlantic climate variability in CCSM3. J. Climate, 19, 2451–2481.
——, R. Tomas, M. Alexander, and D. Lawrence, 2010: The seasonal atmospheric response to projected Arctic sea ice loss in
the late twenty-first century. J. Climate, 23, 333–351.
——, R. Knutti, S. Solomon, and A. S. Phillips, 2012a: Communication of the role of natural variability in future North
American climate. Nat. Climate Change, 2, 775–779.
——, A. Phillips, V. Bourdette, and H. Teng, 2012b: Uncertainty inclimate change projections: The role of internal variability.
Climate Dyn., 38, 527–546.
——, and Coauthors, 2012c: ENSO and Pacific decadal variability
in the Community Climate System Model version 4. J. Climate, 25, 2622–2651.
DeWeaver, E., and C. M. Bitz, 2006: Atmospheric circulation andits effect on Arctic sea ice in CCSM3 simulations at medium
and high resolution. J. Climate, 19, 2415–2436.
——, E. C. Hunke, and M. M. Holland, 2008: Sensitivity of Arctic
sea ice thickness to intermodel variations in the surface energy
budget. Arctic Sea Ice Decline: Observations, Projections,
Mechanisms, and Implications, Geophys. Monogr., Vol. 180,
Amer. Geophys. Union, 77–90.
Dommenget, D., and M. Latif, 2008: Generation of hyper climate
modes. Geophys. Res. Lett., 35, L02706, doi:10.1029/
2007GL031087.
Fetterer, F., K. Knowles, W. Meier, and M. Savoie, 2002: Sea ice
index. National Snow and Ice Data Center, Boulder, CO,digital media. [Available online at http://nsidc.org/data/
g02135.html.]
Fowler, C., 2003: Polar Pathfinder daily 25 km EASE-grid sea ice
motion vectors. National Snow and Ice Data Center, Boulder,CO, digital media. [Available online at http://nsidc.org/data/
docs/daac/nsidc0116_icemotion.gd.html.]
Francis, J. A., and E. Hunter, 2007: Drivers of declining sea ice in
the Arctic winter: A tale of two seas. Geophys. Res. Lett., 34,
L17503, doi:10.1029/2007GL030995.

549

Gent, P. R., and Coauthors, 2011: The Community Climate System
Model version 4. J. Climate, 24, 4973–4991.
Hawkins, E., and R. Sutton, 2009: The potential to narrow uncertainty in regional climate predictions. Bull. Amer. Meteor.
Soc., 90, 1095–1107.
Holland, M. M., and C. M. Bitz, 2003: Polar amplification of climatechange in coupled models. Climate Dyn., 21, 221–232.
——, ——, L.-B. Tremblay, and D. A. Bailey, 2008: The role ofnatural versus forced change in future rapid summer Arctic ice
loss. Arctic Sea Ice Decline: Observations, Projections, Mechanisms, and Implications, Geophys. Monogr., Vol. 180, Amer.
Geophys. Union, 133–150.
Hu, Z.-Z., A. Kumar, B. Jha, and B. Huang, 2012: An analysis offorced and internal variability in a warmer climate in CCSM3.
J. Climate, 25, 2356–2373.
Jahn, A., and Coauthors, 2012: Late-twentieth-century simulationof Arctic sea ice and ocean properties in the CCSM4. J. Climate, 25, 1431–1452.
Kay, J. E., M. M. Holland, and A. Jahn, 2011: Inter-annual to multidecadal Arctic sea ice extent trends in a warming world.
Geophys. Res. Lett., 38, L15708, doi:10.1029/2011GL048008.
——, ——, C. M. Bitz, E. Blanchard-Wrigglesworth, A. Gettelman,
A. Conley, and D. Bailey, 2012: The influence of local feedbacks and northward heat transport on the equilibrium Arctic
climate response to increased greenhouse gas forcing. J. Climate, 25, 5433–5450.
Kvamstø, N. G., P. Skeie, and D. B. Stephenson, 2004: Impact ofLabrador sea-ice extent on the North Atlantic Oscillation. Int.
J. Climatol., 24, 603–612.
Lee, S., T. Gong, N. Johnson, S. B. Feldstein, and D. Pollard, 2011:
On the possible link between tropical convection and theNorthern Hemisphere Arctic surface air temperature change
between 1958 and 2001. J. Climate, 24, 4350–4367.
L’Heureux, M. L., A. Kumar, G. D. Bell, M. S. Halpert, and R. W.
Higgins, 2008: Role of the Pacific-North American (PNA)
pattern in the 2007 Arctic sea ice decline. Geophys. Res. Lett.,
35, L20701, doi:10.1029/2008GL035205.
Li, C., and J. J. Wettstein, 2012: Thermally driven and eddy-driven
jet variability in reanalysis. J. Climate, 25, 1587–1596.
Magnusdottir, G., C. Deser, and R. Saravanan, 2004: The effects ofNorth Atlantic SST and sea ice anomalies on the winter circulation in CCM3. Part I: Main features and storm-track
characteristics of the response. J. Climate, 17, 857–876.
Mahlstein, I., and R. Knutti, 2011: Ocean heat transport as a cause
for model uncertainty in projected Arctic warming. J. Climate,24, 1451–1460.
Maslanik, J. A., C. Fowler, J. Stroeve, S. Drobot, H. J. Zwally,
D. Yi, and W. J. Emery, 2007: A younger, thinner ice cover:
Increased potential for rapid, extensive sea-ice loss. Geophys.
Res. Lett., 34, L24501, doi:10.1029/2007GL032043.
Meehl, G. A., and Coauthors, 2006: Climate change projections forthe twenty-first century and climate change commitment in
the CCSM3. J. Climate, 19, 2597–2616.
——, and Coauthors, 2012: Climate system response to external
forcings and climate change projections in CCSM4. J. Climate,25, 3661–3683.
Moss, R. H., and Coauthors, 2010: The next generation of scenarios
for climate change research and assessment. Nature, 463, 747–
756.
Nghiem, S. V., Y. Chao, G. Neumann, P. Li, D. K. Perovich,
T. Street, and P. Clemente-Colon, 2006: Depletion of perennial sea ice in the east Arctic Ocean. Geophys. Res. Lett., 33,
L17501, doi:10.1029/2006GL027198.

550

JOURNAL OF CLIMATE

North, G. R., T. L. Bell, R. F. Cahalan, and F. J. Moeng, 1982:
Sampling errors in the estimation of empirical orthogonal
functions. Mon. Wea. Rev., 110, 699–706.
Ogi, M., and J. M. Wallace, 2007: Summer minimum Arctic sea ice
extent and the associated summer atmospheric circulation.
Geophys. Res. Lett., 34, L12705, doi:10.1029/2007GL029897.
——, and ——, 2012: The role of summer surface wind anomalies inthe summer Arctic sea ice extent in 2010 and 2011. Geophys.
Res. Lett., 39, L09704, doi:10.1029/2012GL051330.
——, K. Yamazaki, and J. M. Wallace, 2010: Influence of winterand summer surface wind anomalies on summer Arctic sea ice
extent. Geophys. Res. Lett., 37, L07701, doi:10.1029/
2009GL042356.
Overland, J. E., and M. Wang, 2005: The third Arctic climate
pattern: 1930s and early 2000s. Geophys. Res. Lett., 32,
L23808, doi:10.1029/2005GL024254.
——, and ——, 2010: Large-scale atmospheric circulation changes areassociated with the recent loss of Arctic sea ice. Tellus, 62A, 1–9.
Perovich, D. K., B. Light, H. Eicken, K. F. Jones, K. Runciman, andS. V. Nghiem, 2007: Increasing solar heating of the Arctic
Ocean and adjacent seas, 1979–2005: Attribution and role in
the ice-albedo feedback. Geophys. Res. Lett., 34, L19505,
doi:10.1029/2007GL031480.
Polyakov, I. V., and Coauthors, 2005: One more step toward
a warmer Arctic. Geophys. Res. Lett., 32, L17605, doi:10.1029/
2005GL023740.
Rigor, I. G., and J. M. Wallace, 2004: Variations in the age of Arctic
sea-ice and summer sea ice extent. Geophys. Res. Lett., 31,
L09401, doi:10.1029/2004GL019492.
——, ——, and R. L. Colony, 2002: Response of sea ice to the
Arctic Oscillation. J. Climate, 15, 2648–2663.
Rothrock, D. A., Y. Yu, and G. A. Maykut, 1999: Thinning of the
Arctic sea-ice cover. Geophys. Res. Lett., 26 (23), 3469–3472.
Screen, J. A., I. Simmonds, and K. Keay, 2011: Dramatic interannual changes of perennial Arctic sea ice linked to abnormal summer storm activity. J. Geophys. Res., 116, D15105,
doi:10.1029/2011JD015847.
Seierstad, I. A., and J. Bader, 2009: Impact of a projected future Arctic
sea ice reduction on extratropical storminess and the NAO.
Climate Dyn., 33, 937–943, doi:10.1007/s00382-008-0463-x.
Serreze, M. C., and R. G. Barry, 2011: Processes and impacts ofArctic amplification: A research synthesis. Global Planet.
Change, 77, 85–96.
——, M. M. Holland, and J. Stroeve, 2007: Perspectives on the
Arctic’s shrinking sea-ice cover. Science, 315, 1533–1536.
Shimada, K., T. Kamoshida, M. Itoh, S. Nishio, E. Carmack,
F. McLaughlin, S. Zimmerman, and A. Proshutinsky, 2006:
Pacific Ocean inflow: Influence on catastrophic reduction ofsea ice cover in the Arctic Ocean. Geophys. Res. Lett., 33,
L08605, doi:10.1029/2005GL025624.
Simmons, A., S. Uppala, D. P. Dee, and S. Kobayashi, 2007: ERAInterim: New ECMWF reanalysis products from 1989 onwards. ECMWF Newsletter, No. 110, ECMWF, Reading,United Kingdom, 25–35.

VOLUME 27

Smedsrud, L. H., A. Sirevaag, K. Kloster, A. Sorteberg, and
S. Sandven, 2011: Recent wind driven high sea ice area export in the Fram Strait contributes to Arctic sea ice decline.
Cryosphere, 5, 821–829.
Sorteberg, A., T. Furevik, H. Drange, and N. G. Kvamstø, 2005:
Effects of simulated natural variability on Arctic temperature
projections. Geophys. Res. Lett., 32, L18708, doi:10.1029/
2005GL023404.
Stoner, A. M. K., K. Hayhoe, and D. J. Wuebbles, 2009: Assessing
general circulation model simulations of atmospheric teleconnection patterns. J. Climate, 22, 4348–4372.
Stroeve, J. C., M. W. Holland, W. Meier, T. Scambos, and
M. Serreze, 2007: Arctic sea ice decline: Faster than forecast.
Geophys. Res. Lett., 34, L09501, doi:10.1029/2007GL029703.
——, J. Maslanik, M. C. Serreze, I. Rigor, W. Meier, and C. Fowler,
2011: Sea ice response to an extreme negative phase of the
Arctic Oscillation during winter 2009/2010. Geophys. Res.
Lett., 38, L02502, doi:10.1029/2010GL045662.
——, V. Kattsov, A. Barrett, M. Serreze, T. Pavlova, M. Holland,
and W. N. Meier, 2012: Trends in Arctic sea ice extent from
CMIP5, CMIP3 and observations. Geophys. Res. Lett., 39,
L16502, doi:10.1029/2012GL052676.
Tebaldi, C., and R. Knutti, 2007: The use of the multimodel ensemble in probabilistic climate projections. Philos. Trans. Roy.
Soc. London, A365, 2053–2075.
Teng, H., and G. Branstator, 2011: Initial-value predictability ofprominent modes of North Pacific subsurface temperature
in a CGCM. Climate Dyn., 36, 1813–1834, doi:10.1007/
s00382-010-0749-7.
Thompson, D. W. J., and J. M. Wallace, 1998: The Arctic Oscillation signature in the wintertime geopotential height andtemperature fields. Geophys. Res. Lett., 25 (9), 1297–1300.
Tsukernik, M., C. Deser, M. Alexander, and R. Tomas, 2010: Atmospheric forcing of Fram Strait sea ice export: A closer look.
Climate Dyn., 35, 1349–1360, doi:10.1007/s00382-009-0647-z.
Vavrus, S. J., M. M. Holland, A. Jahn, D. A. Bailey, and B. A.
Blazey, 2012: Twenty-first-century Arctic climate change in
CCSM4. J. Climate, 25, 2696–2710.
Vinje, T., 2001: Fram Strait ice fluxes and atmospheric circulation:
1950–2000. J. Climate, 14, 3508–3517.
Wang, J., J. Zhang, E. Watanabe, M. Ikeda, K. Mizobata, J. E.
Walsh, X. Bai, and B. Wu, 2009: Is the dipole anomaly a major
driver to record lows in Arctic summer sea ice extent? Geophys. Res. Lett., 36, L05706, doi:10.1029/2008GL036706.
Wettstein, J. J., and J. M. Wallace, 2010: Observed patterns ofmonth-to-month storm-track variability and their relationshipto the background flow. J. Atmos. Sci., 67, 1420–1437.
Winton, M., 2008: Sea ice-albedo feedback and nonlinear Arctic
climate change. Arctic Sea Ice Decline: Observations, Projections, Mechanisms, and Implications, Geophys. Monogr.,Vol. 180, Amer. Geophys. Union, 111–131.
Wu, B., J. Wang, and J. E. Walsh, 2006: Dipole anomaly in thewinter Arctic atmosphere and its association with sea ice
motion. J. Climate, 19, 210–225.