Home

SearchCollections

Journals

About

Contact us

My IOPscience

Influence of projected Arctic sea ice loss on polar stratospheric ozone and circulation in springThis content has been downloaded from IOPscience. Please scroll down to see the full text.
2014 Environ. Res. Lett. 9 084016(http://iopscience.iop.org/1748-9326/9/8/084016)
View the table of contents for this issue, or go to the journal homepage for more

Download details:
IP Address: 128.117.20.88
This content was downloaded on 01/04/2015 at 21:54

Please note that terms and conditions apply.

Environmental Research Letters
Environ. Res. Lett. 9 (2014) 084016 (9pp)

doi:10.1088/1748-9326/9/8/084016

Inﬂuence of projected Arctic sea ice loss onpolar stratospheric ozone and circulation inspringLantao Sun1, Clara Deser1, Lorenzo Polvani2 and Robert Tomas1
1

National Center for Atmospheric Research, Boulder, CO, USA
Dept of Applied Physics and Applied Mathematics, and Dept of Earth and Environmental Sciences,
Columbia University, New York, USA

2E-mail: lantao@ucar.edu
Received 7 May 2014, revised 30 July 2014
Accepted for publication 31 July 2014
Published 26 August 2014
Abstract

The impact of projected Arctic sea ice loss on the stratosphere is investigated using the Whole
Atmosphere Community Climate Model (WACCM), a state-of-the-art coupled chemistry climate
model. Two 91-year simulations are conducted: one with a repeating seasonal cycle of Arctic sea
ice for the late twentieth-century, taken from the fully coupled WACCM historical run; the otherwith Arctic sea ice for the late twenty-ﬁrst century, obtained from the fully coupled WACCM
RCP8.5 run. In response to Arctic sea ice loss, polar cap stratospheric ozone decreases by 13 DU(34 DU at the North Pole) in spring, conﬁrming the results of Scinocca et al (2009 Geophys. Res.
Lett. 36 L24701). The ozone loss is dynamically initiated in March by a suppression of upwardpropagating planetary waves, possibly related to the destructive interference between the forced
wave number 1 and its climatology. The diminished upward wave propagation, in turn, weakens
the Brewer–Dobson circulation at high latitudes, strengthens the polar vortex, and cools the polar
stratosphere. The ozone reduction persists until the polar vortex breaks down in late spring.
S Online supplementary data available from stacks.iop.org/erl/9/084016/mmedia
Keywords: Arctic sea ice loss, ozone, Brewer–Dobson circulation

1. Introduction(NAM) (Thompson and Wallace 2000) in winter (see Deseret al 2010, Peings and Magnusdottir 2014 and references
therein).
Given the projected impacts of Arctic sea ice loss on the
troposphere, it is reasonable to ask whether the stratospheremight be affected as well. To our knowledge, only one study
has addressed this question using a model with a wellresolved stratosphere, namely Scinocca et al (2009), hereafterreferred to as S09. S09 examined the transient response of the
fully coupled Canadian Middle Atmosphere Model to asudden reduction in sea ice albedo and found that the polar
stratosphere cools in response to Arctic sea ice loss, accompanied by a local reduction in ozone. Other studies based on‘low-top’ atmospheric general circulation models without awell-represented stratosphere also ﬁnd a springtime cooling of
the polar stratosphere in response to present and future Arctic
sea ice loss (Cai et al 2012, Screen et al 2013). In 2011,

Arctic sea ice extent is declining at an accelerating pace, andclimate models project a seasonally ice-free Arctic Ocean by
the middle of this century in response to increasing greenhouse gas (GHG) concentrations (see Stroeve et al 2012, andreferences therein). The sea ice loss is expected to havenumerous consequences for regional climates, including nearsurface warming and increased precipitation at high latitudes(e.g., Screen and Simmonds 2010, Deser et al 2010). It is also
expected to affect the large-scale tropospheric circulation,
manifest as the negative phase of the Northern Annular Mode
Content from this work may be used under the terms of theCreative Commons Attribution 3.0 licence. Any further
distribution of this work must maintain attribution to the author(s) and thetitle of the work, journal citation and DOI.
1748-9326/14/084016+09$33.00

1© 2014 IOP Publishing Ltd

Environ. Res. Lett. 9 (2014) 084016

L Sun et alfollowing several years of below normal ice cover, the Arctic
experienced unprecedented springtime ozone depletion, raising additional concerns about a possible linkage between
Arctic sea ice loss and ozone (Manney et al 2011).
Building upon the results of S09, Cai et al (2012), andScreen et al (2013), we conduct experiments with the Whole
Atmosphere Community Climate Model (WACCM), a stateof-the-art chemistry-climate model with a well-resolved
stratosphere, to study the impact of future Arctic sea ice loss
upon the stratosphere, including its circulation, temperature,and ozone concentration. Unlike S09, who studied the transient response to an abrupt idealized reduction in sea ice in acoupled model context, we here examine the steady-state
atmospheric response to late twenty-ﬁrst century sea ice loss
projected by the fully coupled version of WACCM, driven by
the RCP8.5 GHG scenario. As shown below, future Arctic sea
ice loss suppresses upward-propagating planetary waves inMarch, resulting in a weakened Brewer–Dobson circulation athigh latitudes and an associated reduction in lower stratospheric ozone, which persists until the polar vortex breaks
down in late spring.

2. Model and experimental design
We use WACCM version 4, a high-top model with a horizontal resolution of 1.9° latitude and 2.5° longitude and 66
vertical levels extending from the surface to approximately
140 km. In addition to enhanced vertical resolution in thestratosphere and mesosphere, WACCM incorporates aninteractive stratospheric chemistry package and special
gravity wave parameterizations. These additional features
make this model a better tool for studying the stratospheric
response than more commonly used low-top models. Detailsof the model formulation and WACCMʼs twentieth-century
transient simulation are documented in Marsh et al (2013).
To address the impact of future sea ice loss upon thestratosphere, we have conducted a pair of 91-year experiments. The control experiment is speciﬁed with a repeating
seasonal cycle of sea ice concentration (SIC) and sea surface
temperature (SST) averaged over the period 1980–1999,
obtained from the average of a three-member ensemble oftwentieth-century simulations with the corresponding fully
coupled version of WACCM (Marsh et al 2013). The perturbation experiment is speciﬁed with a repeating seasonal
cycle of Arctic SIC averaged over the period 2080–2099,
obtained from the single available twenty-ﬁrst century simulation of the fully coupled version of WACCM, forced by‘Representative Concentration Pathway (RCP) 8.5’ (corresponding to a radiative forcing level of approximately 8.5 W
m−2 by 2100). In both experiments, the radiative forcings andozone-depleting substances (ODS) are ﬁxed at the year 2000
so as to isolate the impact of Arctic sea ice loss. Note that we
do not change SIC in the Antarctic. Following convention,
sea ice thickness is speciﬁed to be 2 m in the Arctic and 1 m in
the Antarctic.

Figure 1. (a) Average Arctic sea ice concentration (SIC; %)simulated by WACCM fully coupled (left) historical simulation
during 1980–1999, (middle) RCP8.5 simulation during 2080–2099,
and (right) their difference (RCP8.5 minus historical) in (top) Marchand (bottom) September. (b) Monthly changes of Sea Ice Extent(SIE; gray bars; 106 km2) between the RCP8.5 and historicalsimulations and the monthly responses of the net surface energy ﬂux(blue curve), sensible plus latent heat energy ﬂux (yellow curve), andlong-wave radiative ﬂux (green curve) averaged over the Arctic
ocean (W m−2, positive values upward).

The SSTs in the perturbation experiment are the same as
those in the control experiment, except in areas where thefractional ice cover in the late twenty-ﬁrst century is less thanthat in the late twentieth-century. In these cases, the SSTs are
set to their 2080–2099 values from WACCM RCP8.5 run.
This approach takes into account not only the sea ice loss, butalso the local warming of the sea surface that occurs inassociation with the ice loss (see also Screen et al 2013). It
should be noted, however, that inclusion of the local SST
warming has only a small (∼15%) effect on the magnitude of
the Arctic Oceanʼs net surface heat ﬂux response in similar
experiments with Community Atmosphere Model Version 4(CAM4) (Deser et al 2014). Differencing the perturbation andcontrol experiments yields the impact of future Arctic sea ice
loss and associated local SST warming on the atmosphere. A
Studentʼs t-test is used to estimate the statistical signiﬁcanceof the responses. We discard the ﬁrst year of each 91-year
simulation from our analysis.
Figure 1(a) shows the prescribed SIC distribution in
March (time of maximum ice extent) and September (time of2

Environ. Res. Lett. 9 (2014) 084016

L Sun et al

minimum ice extent) for the control and perturbation
experiments and their difference. WACCMʼs simulation of
the present-day sea ice distribution is in good agreement withobservations (e.g., Stroeve et al 2012), although there is someoverestimation of the concentrations within the marginal ice
zones. In both months, the sea ice cover contracts poleward in
the late twenty-ﬁrst century compared to the late twentiethcentury. In March, the ice loss is limited to the marginal ice
zones in the North Atlantic and Paciﬁc. In September, however, ice loss occurs throughout the Arctic Ocean, and muchof the basin is nearly ice-free by the end of the twenty-ﬁrst
century.
The differences in the seasonal cycles of sea ice extent(SIE) between the perturbation and control experiments areshown in ﬁgure 1(b), together with the monthly responses ofthe surface energy ﬂuxes (positive upward) averaged over theArctic Ocean. The largest ice loss between the late twentieth(1980–1999) and late twenty-ﬁrst (2080–2099) centuries
occurs in August–October (∼6×106 km2; gray bars), with arelative small seasonal dependence to the accompanying local
SST increase (minimum values ∼0.3°C in February–April
and maximum values ∼ 1.4°C in July–September; notshown). By contrast, the net surface energy ﬂux response(blue curve) exhibits a pronounced seasonal cycle, withmaximum values in winter (∼40 W m−2 in December–January) and minimum values in summer (∼5 W m−2 inJune–August), consistent with previous studies (e.g., Deseret al 2010, Peings and Magnusdottir 2014). The net surface
energy ﬂux is dominated by the turbulent (sensible plus
latent) component (yellow curve). Since sea ice affects theatmosphere via the surface energy ﬂuxes, the delay between
the maximum ice loss and the peak surface energy ﬂux
response has implications for the timing of the atmosphericcirculation response (Deser et al 2010). Speciﬁcally, theimpact of Arctic sea loss on the atmosphere is largest in thewinter even though ice loss peaks in the fall.

Figure 2. (a) Distribution of zonally averaged total column ozone(DU) as a function of month and latitude for the control run
climatology (contours; contour interval of 25 DU) and the response
to Arctic sea ice loss (shading; stippled region indicates >95%
statistical signiﬁcance based on a two-sided Studentʼs t-test). (b)Time series of total column ozone (DU) averaged over the polar cap(60 °N–90°N) in March from the control (blue) and perturbation(red) experiments. The horizontal dashed lines show the time-meanfor each experiment.

with an overall reduction of 13 DU. However, there is large
interannual variability in both experiments, with an approximate range of 150 DU. This highlights the importance oflong simulations to establish the statistical signiﬁcance of the
response in the presence of internal variability.

3.2. Mechanism of March ozone reduction

To investigate the mechanism of the March ozone reduction,
we show the responses (shading) and control run climatologies (contours) of ozone mixing ratio and residual mean circulation in ﬁgure 3(a) and (b), respectively. The residual
mean circulation approximates the mean Lagrangian transportof mass through the stratosphere and is known as the Brewer–Dobson circulation (BDC). The BDC climatology ischaracterized by tropical upwelling and extratropical downwelling and is responsible for transporting ozone from theequatorial upper stratosphere (where ozone is produced) tothe high-latitude lower stratosphere (e.g., Shepherd 2008).
The circulation response to Arctic sea ice loss shows aweakening of the poleward and downwelling branches of the
BDC at high latitudes. These BDC changes result in diminished transport of ozone into the polar cap in the lower stratosphere, qualitatively consistent with the reduction of theozone mixing ratio in this region. We also analyze lower
stratospheric changes in the area with a temperature below
196 K, the threshold for the formation of nitric acid trihydrate
clouds (Hanson and Mauersberger 1988), and ﬁnd a notable
increase in response to sea ice loss. While a detailed ozone
budget is beyond the scope of this short paper, this suggests

3. Results
3.1. Ozone response to Arctic sea ice loss

Figure 2(a) shows the total column ozone response to projected sea ice loss as a function of latitude and month(shading), superimposed upon the control run climatology(contours). The Arctic ozone climatology in WACCM ischaracterized by a distinct seasonal cycle, with a maximum(∼400–425 DU) in spring and a minimum (∼300–325 DU) in
fall, consistent with observations. In response to sea ice loss, a
substantial and statistically signiﬁcant polar ozone reduction
is found in spring (March–April), with maximum amplitudeof 34 DU at the North Pole. To illustrate the magnitude of the
response relative to the internal variability, we show therecords of March polar cap ozone for each year of the control
and perturbation simulations (ﬁgure 2(b)). Consistent withﬁgure 2(a), March polar ozone is reduced in most years of theperturbation experiment compared to the control experiment,
3

Environ. Res. Lett. 9 (2014) 084016

L Sun et al

Figure 3. Distributions of March zonally averaged ozone and atmospheric circulation as a function of pressure and latitude for the control run
climatology (contours) and the response to Arctic sea ice loss (shading; stippling indicates the 95% signiﬁcance level). (a) Ozone mixing ratio(ppm), (b) Brewer–Dobson circulation (109 kg s−1), (c) Eliassen–Palm (E–P) ﬂux vector and its divergence (m s−1 day−1; shading). Note thatthe climatological E–P ﬂuxes and divergence ﬁelds are not shown. (d) Zonal wind (m s−1), (e) Temperature (K). The contour intervals forozone mixing ratio, Brewer–Dobson circulation, zonal wind, and temperature are 1 ppm, 8 × 108 kg s−1, 5 m s−1, and 10 K, respectively. For
better display, we plot only those E–P vectors whose vertical component exceeds 10 Pa m s−2 in absolute value. Note that a nonlinear color
scale is used for the zonal wind response (interval of 0.25 m s−1 for absolute value less than 1 m s−1 and interval of 1 m s−1 for absolute value
larger than 1 m s−1).

wave drivings in the stratosphere. ﬁgure 3(c) shows the
response of the Eliassen–Palm vector and its divergence, ameasure of resolved Rossby wave propagation and dissipation, in March. The climatology is characterized by upward
wave propagation (upward Eliassen–Palm (E–P) vector) andwave dissipation (E–P convergence) in the extratropicalstratosphere (not shown). In response to sea ice loss, there is a
downward anomaly in the E–P vectors, indicating lessthat both chemistry and transport play a role in the ozone loss
found in our experiments.
Previous studies (e.g., Ueyama and Wallace 2010, Chen
and Sun 2011, Gerber 2012) have suggested that the BDC ismainly driven by stratospheric extratropical planetary waves
and gravity waves (Butchart et al 2010a). Therefore we
examine whether the weakening of the BDC in response to
sea ice loss might be related to planetary wave and/or gravity
4Environ. Res. Lett. 9 (2014) 084016

L Sun et al

Figure 4. Forced (shading) and climatological (contour) stationary wave number (top) one and (bottom) two geopotential height ﬁelds at (a)60°N and (b) 100 hPa in March. Results are shown as longitude-pressure (hPa) cross-sections at 60°N in the left panels and longitude-latitude
maps at 100 hPa in the right panels. The contour intervals are 30 m for cross-sections at 60°N and 10 m for maps at 100 hPa. The weighted
spatial correlation cofﬁcients (R) between the contour and shading is given in each panel.

upward wave propagation into the stratosphere. Associatedwith this, there is reduced wave dissipation (positive E–P
divergence anomaly) in the extratropical stratosphere. In
order to establish the connection between changes in resolved
and parameterized wave drag and the BDC more quantitatively, we use the downward control principle (Haynes
et al 1991) to diagnose the extratropical residual streamfunction at 70 hPa associated with the stratospheric wave dragchange. The results indicate that in our model almost all of the
BDC reduction within the polar cap is caused by the resolved
planetary waves in the stratosphere, with gravity wave dragmaking only a minor contribution (Supplementary Materials,
available at stacks.iop.org/erl/9/084016/mmedia).
One may wonder why the upward wave propagation into
the stratosphere is suppressed in March. Recent studies havesuggested that forced planetary waves could interfere with theclimatology so as to suppress or enhance the upward wave
propagation (Garﬁnkel et al 2010, Fletcher and Kushner 2011, Smith and Kushner 2012). Figure 4 shows theforced (contour) and climatological (shading) stationary wave
numbers 1 and 2 geopotential height ﬁelds at 60°N and 100
hPa in March. Both the forced and climatological wave 1
exhibit a westward tilt with height, corresponding to an
upward propagating Rossby wave. At 100 hPa, the wave 1
response is largely out of phase with the control run climatology. The weighted spatial correlations at 60°N and 100 hPa
are -0.40 and -0.54, respectively. This suggests that, qualitatively, the forced wave 1 is able to destructively interfere withthe climatology, resulting in a suppression of upward wave
propagation. In contrast, the wave 2 response is more in phasewith the climatology below 10 hPa, with the weighted spatial
correlations of 0.42 at 60°N and 0.82 at 100 hPa between thetwo ﬁelds. Therefore, wave 2 interference does not appear to
explain the suppression of upward wave propagation. Further,
the amplitude of the response is considerably larger for wave
1 than wave 2 (96 m compared to 7 m at 60°N and 10 hPa),
another indication that wave 1 is more important than wave 2
5Environ. Res. Lett. 9 (2014) 084016

L Sun et alin suppressing upward wave propagation via linear wave
interference. Moreover, the E–P ﬂux diagnostics for eachwave number separately conﬁrms that the suppression of
upward wave propagation and wave dissipation in the stratosphere is indeed dominated by wave number 1 (not shown).
Collectively these diagnostics point to the role of wave 1 incausing the stratospheric circulation response in March.

mean zonal wind at 10 hPa and 70°N (jet core latitude) drops
below zero without returning above 10 m s−1 until the following autumn. Figure 5(a) shows the time series of polar
vortex breakdown dates in both experiments. The mean datein the control experiment is April 24, approximately three
weeks later than in the NCEP/NCAR Reanalysis (Blacket al 2006). This is a well-known bias of most current-generation chemistry-climate models (Butchart et al 2010b). In
response to sea ice loss, the mean breakdown date is delayedby an average of six days (statistically signiﬁcant from zero atthe 97% conﬁdence level), with the largest effect on the earlyside of the breakdown date distribution.
The timing of polar vortex breakdown is closely connected to the evolution of the ozone response at high latitudes.
As the vortex breaks down in late April in the perturbation
experiment (recall in ﬁgure 5(a) that the mean breakdown
date is April 30), strong isentropic mixing brings high-ozone
air from mid-latitudes into the polar region, abruptly terminating polar ozone depletion. This explains the timing of theozone response to Arctic sea ice loss in ﬁgure 2(a), which
emerges in March and persists through April until the polar
vortex is ﬁnally destroyed.
To further explore how the circulation and ozone
responses vary with the timing of the breakdown of the polar
vortex, we select the ten earliest and ten latest vortex breakdown years from the control and perturbation experiments
and construct composites for 10-hPa zonal wind and total
column ozone (ﬁgures 5(b) and (c)). The polar vortex isstrengthened in March for both the early and late breakdown
composites in response to sea ice loss, with somewhat larger
magnitude for the early breakdown years. The response in thelate breakdown composite, however, persists until the middleof May, about 45 days longer than for the early breakdown
composites. The ozone climatology for the early breakdown
years shows larger values at high latitudes in spring than thatfor the late breakdown years, consistent with observations(Strahan et al 2013) and earlier modeling studies (Salby andCallaghan 2007). The polar ozone response to Arctic sea ice
loss follows the circulation response: a larger loss for the
early breakdown years that disappears quickly in April (leftpanel of ﬁgure 5(b)). In contrast, ozone loss for the latebreakdown years is more persistent, lasting until May.

3.3. Zonal wind and temperature responses

The zonal-mean responses of zonal wind and temperature in
March as a function of latitude and height are shown inﬁgures 3(d) and (e), respectively. In the troposphere, the zonal
wind response is characterized by a statistically signiﬁcant
meridional dipole in the middle latitudes, with reduced westerlies in the latitude band 45–70°N and enhanced westerlies
in the band 20–40°N. This structure implies an equatorward
shift of the midlatitude jet and projects onto the negative
phase of the NAM, consistent with many other studies basedon both comprehensive and idealized models (e.g., Deseret al 2010, Butler et al 2010, Peings and Magnusdottir 2014).
In the stratosphere, there is a marked increase of thezonal winds at high latitudes, with values exceeding 6 m s−1,
and a smaller amplitude reduction of the zonal winds in
middle latitudes (ﬁgure 3(d)). The strengthening of the polar
vortex is to be expected, given the reduced wave dissipationin the stratosphere, as the latter is responsible for decelerating
the zonal wind. From the thermal wind relation, a stronger
polar vortex is associated with a cooler polar lower stratosphere as seen in ﬁgure 3(e), although ozone reduction may
also contribute to the cooling of the polar stratosphere as aresult of diminished absorption of solar ultraviolet radiation.
Near the surface, large warming in the polar region(ﬁgure 3(e)) is a result of an increase of the upward turbulent
heat ﬂux as a consequence of sea ice loss.
3.4. The role of stratospheric polar vortex final breakdown inozone response

So far we have focused on the role of the residual circulation
as a mechanism to explain the reduction of March ozone in
response to projected Arctic sea ice loss. However, the mean
meridional circulation is not the only way to transport ozone
into the polar cap. Isentropic mixing across the vortex edge
can also bring ozone-rich mid-latitude air into the polar
region. This quasi-horizontal mixing weakens in response to
Arctic sea ice loss due to reduced upward planetary-wave
propagation and wave breaking in the stratosphere (e.g.,
Shepherd 2008) and may partly contribute to the polar ozone
loss in March.
The timing of the breakdown of the polar vortex, also
known as the stratospheric ‘ﬁnal warming’, is an indicator ofstrong isentropic mixing and has been suggested to affectozone at interannual timescales in the Northern Hemisphere(Salby and Callaghan 2007, Strahan et al 2013). Therefore,
we examine the timing of polar vortex breakdown in bothexperiments. Following Black et al (2006), we deﬁne theonset date as the last day when the 5-day running mean zonal-

4. Summary and discussion
Our WACCM experiments, contrasting prescribed presentday and future sea ice conditions, indicate that late twentyﬁrst century Arctic sea ice loss, if taken in isolation from otherprojected changes (e.g., GHGs and ODS), will induce a signiﬁcant reduction in stratospheric polar ozone in association
with a weakening of the BDC and a strengthening of thestratospheric polar vortex in spring. The March reduction ofpolar ozone in response to Arctic sea ice loss is initiateddynamically via a suppression of upward-propagating planetary waves, possibly related to the destructive interference
between the forced wave number 1 and its climatology. The
6

Environ. Res. Lett. 9 (2014) 084016

L Sun et al

Figure 5. (a) Time series of polar vortex breakdown dates in the control (blue) and perturbation (red) experiments. Dashed lines show themean onset dates for each experiment. Circle and rectangle markers indicate the ten earliest and ten latest polar vortex breakdown years,
respectively. (b) Composites of climatological zonal wind at 10 hPa (contour) and its response to sea ice loss (shading) based on the ten
earliest (denoted Early) and ten latest (denoted Late) polar vortex breakdown years. (c) As in (b), but for total column ozone. The contour
intervals for zonal wind and total column ozone are 5 m s−1 and 25 DU, respectively.

reduced wave dissipation in the extratropical stratosphere, inturn, induces a weakening of the BDC, strengthening of thepolar vortex, and cooling of the polar stratosphere. The ozone
reduction persists through April until the polar vortex
breaks down.
Our results are in qualitative agreement with S09,
although our experimental design differs. In particular, S09
imposed a sudden reduction in sea ice albedo in their fully
coupled model to represent the maximum possible forcing inwhich all sea ice is melted in summer, whereas we haveprescribed sea ice conditions taken from WACCMʼs fully
coupled response to historical and RCP8.5 radiative forcing to
the uncoupled WACCM atmospheric model. Therefore, themagnitude of our sea ice forcing is probably more realistic in

the sense that it is taken from a coupled modelʼs simulation of
future climate. In addition, the seasonality of our sea ice loss
is probably more realistic than in S09 for the same reason. In
particular, sea ice albedo methodology underestimates sea ice
loss in winter (Deser et al 2014), the season in which the
projected sea ice melt is expected to have the largest impact
on surface energy ﬂuxes and atmospheric circulation. Moreover, the response in S09ʼs experiments is transient and only
appears one decade after the ice albedo change is applied. In
our experiments, however, since prescribed sea ice conditions
are repeated for each year of the 91-year runs, we obtain thesteady-state atmospheric response to sea ice loss. Despite
these differences, the agreement of the response acrossmodels highlights the robustness of the results.
7

Environ. Res. Lett. 9 (2014) 084016

L Sun et al

Acknowledgments

To further document the robustness of our results, wehave conducted analogous experiments with CAM4, a ‘lowtop’ model without interactive chemistry, with horizontal
resolution of 0.9° latitude and 1.25° longitude and 26 verticallevels (eight levels above 100 hPa). The prescribed sea ice
loss in the CAM4 experiments (taken from the coupled model
Community Climate System Model version 4 (CCSM4)) issimilar to that in WACCM (not shown). The comparison of
the atmospheric circulation response to Arctic sea ice loss
between WACCM and CAM4 is shown in ﬁgure S2 in the
Supplementary Materials. The response in CAM4 bears a
striking resemblance to the one in WACCM, including
strengthening of the polar vortex and polar cap cooling in thestratosphere. However, the magnitudes are approximately
60% weaker, and the statistical signiﬁcance is a little lower(around 90%) in CAM4 than in WACCM. The weaker
response in CAM4 may be due to several factors, including
differences in model physics, the lack of interactive stratospheric chemistry, and the lack of a well-resolved stratosphere. We note that the climatological polar vortex is muchstronger in CAM4 than in WACCM (maximum zonal-mean
zonal wind at 10 hPa and 70°N is 45 m s−1 in CAM4 compared to 25 m s−1 in WACCM, likely due to the absence ofstratospheric sudden warmings in CAM4), a factor whichmight also affect the response to sea ice forcing (e.g., Peingsand Magnusdottir 2014). Nevertheless, the qualitative
resemblance between the responses in CAM4 and WACCM
adds robustness to the ﬁndings reported here.
It should be noted that the March zonal wind changes
caused by the loss of Arctic ice have opposite signs across thetropopause, with a negative NAM response in the troposphere(i.e., an equatorward shift of the eddy-driven jet) accompanying a positive NAM response in the stratosphere (i.e., astronger polar vortex). This is in contrast to the zonal windchanges in winter, which show negative NAM responses inboth the troposphere and stratosphere (i.e., an equatorward
shift of the eddy-driven jet with a weakened polar vortex; notshown). The winter response more closely resembles the
typical pattern of stratosphere-troposphere coupling in which
NAM anomalies have the same sign across the tropopause(e.g., Gerber and Polvani 2009). The dynamical mechanisms
behind the different vertical structures of the circulation
responses in winter and spring warrant further investigation.
Finally, we note that in our experiments there is a substantial reduction of polar ozone in both March and April (13and 14 DU, respectively) in response to Arctic sea ice loss.
Given that ozone recovery proceeds at an approximately
linear pace of 11 DU decade−1 in the fully coupled WACCM
RCP8.5 simulation (not shown), this means that Arctic sea ice
loss acts to delay ozone recovery by approximately 12 years,
similar to the delay estimated in S09. However, since theNorthern Hemisphere polar ozone is expected to recover by
the middle of this century (earlier than the future sea ice
forcing used here), this delay time is likely to be an upperbound.

The authors would like to thank Drs. Walter Robinson, Paul
Kushner, Gang Chen, and Yaga Richter for valuable discussions, and Adam Phillips for help with NCL graphics. We
also appreciate the constructive comments from three anonymous reviewers. This study is supported by the National
Science Foundation (NSF) Arctic System Science Program.
NCAR is sponsored by the NSF. The work of LMP is funded,in part, by a grant from the NSF.

References
Black R X, McDaniel B A and Robinson W A 2006 Stratospheretroposphere coupling during spring onset J. Clim. 19 4891–901
Butchart N et al 2010a Chemistry-climate model simulations oftwenty-ﬁrst century stratospheric climate and circulation
changes J. Clim. 23 5349–74
Butchart N et al 2010b in SPARC report on the evaluation ofchemistry-climate models ed V Eyring, T G Shepherd, D W
Waugh SPARC Report No. 5, WCRP-132, WMO/TD-1526
Butler A H, Thompson D W J and Heikes R 2010 The steady-state
atmospheric circulation response to climate change-likethermal forcings in a simple general circulation model J. Clim.
23 3474–96
Cai D, Dameris M, Garny H and Runde T 2012 Implications of allseason Arctic sea ice anomalies on the stratosphere Atmos.
Chem. Phys. 12 11819–31
Chen G and Sun L 2011 Mechanisms of the tropical upwellingbranch of the Brewer–Dobson circulation: the role ofextratropical waves J. Atmos. Sci. 68 2878–92
Deser C, Tomas R and Sun L 2014 The role of ocean-atmosphere
coupling in the zonal-mean atmospheric response to Arctic sea
ice loss J. Clim. submittedDeser C, Tomas R A, Alexander M A and Lawrence D 2010 The
seasonal atmospheric response to projected Arctic sea ice loss
in the late twenty-ﬁrst century J. Clim. 23 333–51
Fletcher C G and Kushner P J 2011 The role of linear interference inthe annular mode response to tropical SST forcing J. Clim.
24 778Garﬁnkel C I, Hartmann D L and Sassi F 2010 Troposphericprecursors of anomalous northern hemisphere stratosphericpolar vortices J. Clim. 23 3282
Gerber E P 2012 Stratospheric versus tropospheric control of thestrength and structure of the Brewer–Dobson circulation
J. Atmos. Sci. 69 2857–77
Gerber E P and Polvani L M 2009 Stratosphere-troposphere
coupling in a relatively simple AGCM: the importance ofstratospheric variability J. Clim. 22 1920–33
Hanson D and Mauersberger K 1988 Laboratory studies of the nitricacid trihydrate: implications for the south polar stratosphere
Geophys. Res. Lett. 15 855–8
Haynes P H, Mcintyre M E, Shepherd T G, Marks C J and Shine K P
1991 On the ‘downward control’ of extratropical diabatic
circulations by eddy-induced mean zonal forces. J. Atmos. Sci.
48 651–78
Manney G L et al 2011 Unprecedented Arctic ozone loss in 2011
Nat. Geosci. 478 469–75
Marsh D R, Mills M J, Kinnison D E, Lamarque J-F, Calvo N and
Polvani L M 2013 Climate change from 1850 to 2005
simulated in CESM1(WACCM) J. Clim. 26 7372–91
Peings Y and Magnusdottir G 2014 Response of the Northern
Hemisphere atmospheric circulation to current and projected
Arctic sea ice decline : a numerical study with CAM5 J. Clim.
27 244–64
8

Environ. Res. Lett. 9 (2014) 084016

L Sun et alStrahan S E, Douglass A R and Newman P A 2013 The
contributions of chemistry and transport to low arctic ozone inmarch 2011 derived from Aura MLS observations J. Geophys.
Res.: Atmos. 118 1563–76
Stroeve J C, Kattsov V, Barrett A, Serreze M, Pavlova T,
Holland M and Meier W N 2012 Trends in Arctic sea ice
extent from CMIP5, CMIP3 and observations Geophys. Res.
Lett. 39 L165020
Stroeve J C, Serreze M C, Holland M M, Kay J E, Malanik J and
Barrett A P 2012 b: The Arctic rapidly shrinking sea ice cover:a research synthesis Clim. Change 110 1005–27
Thompson D W J and Wallace J M 2000 Annular modes in theextratropical circulation. Part I: month-to-month variability
J. Clim. 13 1000–16
Ueyama R and Wallace J M 2010 To what extent does high-latitude
wave forcing drive tropical upwelling in the Brewer–Dobson
circulation? J. Atmos. Sci. 67 1232–46

Salby M L and Callaghan P F 2007 Inﬂuence of planetary wave
activity on the stratospheric ﬁnal warming and spring ozone
J. Geophys. Res. 112 20111
Scinocca J F, Reader M C, Plummer D A, Sigmond M, Kushner P J,
Shepherd T G and Ravishankara A R 2009 Impact of sudden
Arctic sea ice loss on stratospheric polar ozone recovery
Geophys. Res. Lett. 36 L24701
Screen J a and Simmonds I 2010 The central role of diminishing sea
ice in recent Arctic temperature ampliﬁcation Nature 464
1334–7
Screen J A, Simmonds I, Deser C and Tomas R A 2013 The
atmospheric response to three decades of observed arctic sea
ice loss J. Clim. 26 1230–48
Shepherd T G 2008 Dynamics, stratospheric ozone, and climatechange Atmos. Ocean 46 117–38
Smith K L and Kushner P J 2012 Linear interference and theinitiation of extratropical stratosphere-troposphere interactions
J. Geophys. Res. 117 D13107

9