Unprecedented Arctic ozone loss in 2011

Abstract

Chemical ozone destruction occurs over both polar regions in local winter-spring. In the Antarctic, essentially complete removal of lower-stratospheric ozone currently results in an ozone hole every year, whereas in the Arctic, ozone loss is highly variable and has until now been much more limited. Here we demonstrate that chemical ozone destruction over the Arctic in early 2011 was--for the first time in the observational record--comparable to that in the Antarctic ozone hole. Unusually long-lasting cold conditions in the Arctic lower stratosphere led to persistent enhancement in ozone-destroying forms of chlorine and to unprecedented ozone loss, which exceeded 80 per cent over 18-20 kilometres altitude. Our results show that Arctic ozone holes are possible even with temperatures much milder than those in the Antarctic. We cannot at present predict when such severe Arctic ozone depletion may be matched or exceeded.

Full text

ARTICLE
doi:10.1038/nature10556
Unprecedented Arctic ozone loss in 2011
Gloria L. Manney
1,2
, Michelle L. Santee
1
, Markus Rex
3
, Nathaniel J. Livesey
1
, Michael C. Pitts
4
, Pepijn Veefkind
5,6
, Eric R. Nash
7
,
Ingo Wohltmann
3
, Ralph Lehmann
3
, Lucien Froidevaux
1
, Lamont R. Poole
8
, Mark R. Schoeberl
9
, David P. Haffner
7
,
Jonathan Davies
10
, Valery Dorokhov
11
, Hartwig Gernandt
3
, Bryan Johnson
12
, Rigel Kivi
13
, Esko Kyro
¨
13
, Niels Larsen
14
,
Pieternel F. Levelt
5,6,15
, Alexander Makshtas
16
, C. Thomas McElroy
10
, Hideaki Nakajima
17
, Maria Concepcio
´
n Parrondo
18
,
David W. Tarasick
10
, Peter von der Gathen
3
, Kaley A. Walker
19
& Nikita S. Zinoviev
16
Chemical ozone destruction occurs over both polar regions in local winter–spring. In the Antarctic, essentially complete
removal of lower-stratospheric ozone currently results in an ozone hole every year, whereas in the Arctic, ozone loss is
highly variable and has until now been much more limited. Here we demonstrate that chemical ozone destruction over
the Arctic in early 2011 was—for the first time in the observational record—comparable to that in the Antarctic ozone
hole. Unusually long-lasting cold conditions in the Arctic lower stratosphere led to persistent enhancement in
ozone-destroying forms of chlorine and to unprecedented ozone loss, which exceeded 80 per cent over 18–20
kilometres altitude. Our results show that Arctic ozone holes are possible even with temperatures much milder than
those in the Antarctic. We cannot at present predict when such severe Arctic ozone depletion may be matched or
exceeded.
Since the emergence of the Antarctic ‘ozone hole’ in the 1980s
1
and
elucidation of the chemical mechanisms
2–5
and meteorological con-
ditions
6
involved in its formation, the likelihood of extreme ozone
depletion over the Arctic has been debated. Similar processes are at
work in the polar lower stratosphere in both hemispheres, but differ-
ences in the evolution of the winter polar vortex and associated polar
temperatures have in the past led to vastly disparate degrees of spring-
time ozone destruction in the Arctic and Antarctic. We show that
chemical ozone loss in spring 2011 far exceeded any previously
observed over the Arctic. For the first time, sufficient loss occurred
to reasonably be described as an Arctic ozone hole.
Arctic polar processing in 2010–11
In the winter polar lower stratosphere, low temperatures induce
condensation of water vapour and nitric acid (HNO
3
) into polar
stratospheric clouds (PSCs). PSCs and other cold aerosols provide
surfaces for heterogeneous conversion of chlorine from longer-lived
reservoir species, such as chlorine nitrate (ClONO
2
) and hydrogen
chloride (HCl), into reactive (ozone-destroying) forms, with chlorine
monoxide (ClO) predominant in daylight
5,7
.
In the Antarctic, enhanced ClO is usually present for 4–5 months
(through to the end of September)
8–11
, leading to destruction of most
of the ozone in the polar vortex between ,14 and 20 km altitude
7
.
Although ClO enhancement comparable to that in the Antarctic
occurs at some times and altitudes in most Arctic winters
9
, it rarely
persists for more than 2–3 months, even in the coldest years
10
. Thus
chemical ozone loss in the Arctic has until now been limited, with
largest previous losses observed in 2005, 2000 and 1996
7,12–14
.
The 2010–11 Arctic winter–spring was characterized by an
anomalously strong stratospheric polar vortex and an atypically long
continuously cold period. In February–March 2011, the barrier to
transport at the Arctic vortex edge was the strongest in either hemi-
sphere in the last ,30 years (Fig. 1a, Supplementary Discussion).
The persistence of a strong, cold vortex from December through to
the end of March was unprecedented. In the previous years with most
ozone loss, temperatures (T) rose above the threshold associated with
chlorine activation (T
act
, near 196 K, roughly the threshold for the
potential existence of PSCs) by early March (Fig. 1b, Supplementary
Figs 1, 2). Only in 2011 and 1997 have Arctic temperatures below T
act
persisted through to the end of March, sporadically approaching a
vortex volume fraction similar in size to that in some Antarctic winters
(Fig. 1b). In 1996–97, however, the cold volume remained very limited
until mid-January and was smaller than that in 2011 at most times
during late January through to the end of March (Fig. 1b, Supplemen-
tary Figs 1, 2).
Daily minimum temperatures in the 2010–11 Arctic winter were
not unusually low, but the persistently cold region was remarkably
deep (Supplementary Figs 1, 2). Temperatures were below T
act
for
more than 100 days over an altitude range of ,15–23 km, compared
to a similarly prolonged cold period over only ,20–23 km altitude in
1997; below ,19 km altitude, T , T
act
continued for ,30 days longer
in 2011 than in 1997 (Supplementary Fig. 1b). In 2005, the previous
year with largest Arctic ozone loss
7
, T , T
act
occurred for more than
100 days over ,17–23 km altitude, but all before early March.
The winter mean volume of air in which PSCs may form (that is,
with T , T
act
), V
psc
, is closely correlated with the potential for ozone
loss
7,15–17
. In 2011, V
psc
(as a fraction of the vortex volume) was the
largest on record (Fig. 1c). Both large V
psc
and cold lingering well into
spring are important in producing severe chemical loss
7,15,16
, and
2010–11 was the only Arctic winter during which both conditions
have been met. Much lower fractional V
psc
in 1997 than in 1996, 2000,
2005 or 2011 (Fig. 1c) is consistent with less ozone loss that year
16,17
.
1
Jet Propulsion Laboratory, California Institute of Technology, Pasadena, California 91109, USA.
2
New Mexico Institute of Mining and Technology, Socorro, New Mexico 87801, USA.
3
Alfred Wegener
Institute for Polar and Marine Research, D-14473 Potsdam, Germany.
4
NASA Langley Research Center, Hampton, Virginia 23681, USA.
5
Royal Netherlands Meteorological Institute, 3730 AE De Bilt, The
Netherlands.
6
Delft University of Technology, 2600 GA Delft, The Netherlands.
7
Science Systems and Applications, Inc., Lanham, Maryland 20706, USA.
8
Science Systems and Applications, Inc., Hampton,
Virginia 23666, USA.
9
Science and Technology Corporation, Lanham, Maryland 20706, USA.
10
Environment Canada, Toronto, Ontario, Canada M3H 5T4.
11
Central Aerological Observatory, Dolgoprudny
141700, Russia.
12
NOAA Earth System Research Laboratory, Boulder, Colorado 80305, USA.
13
Arctic Research Center, Finnish Meteorological Institute, 99600 Sodankyla
¨
, Finland.
14
Danish Climate
Center, Danish Meteorological Institute, DK-2100 Copenhagen, Denmark.
15
Eindhoven University of Technology, 5600 MB Eindhoven, The Netherlands.
16
Arctic and Antarctic Research Institute, St
Petersburg 199397, Russia.
17
National Institute for Environmental Studies, Tsukuba-city, 305-8506, Japan.
18
National Institute for Aerospace Technology, 28850 Torrejo
´
n De Ardoz, Spain.
19
University of
Toronto, Toronto, Ontario, Canada M5S 1A7.
00 MONTH 2011 | VOL 000 | NATURE | 1
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Factors playing secondary parts in governing interannual variability
in ozone destruction, including vortex strength, structure and posi-
tion relative to the cold region, also favour large loss in 2011 (Sup-
plementary Figs 2, 3, Supplementary Discussion). However, despite
the fraction of the vortex with T , T
act
and mid-March temperatures
sporadically approaching those seen in the Antarctic (Fig. 1b,
Supplementary Fig. 1a), even in 2011 temperatures were much higher,
and the cold regions much smaller, than those in most Antarctic
winters.
Satellite trace-gas and PSC measurements highlight the stark con-
trast between polar processing in 2010–11 and that in typical Arctic
winters, and the parallels with Antarctic conditions (Figs 2, 3). In 2011,
PSCs or aerosols were abundant until mid-March (Fig. 3a; consistent
with a deep region with T , T
act
, Fig. 3b), much later than usual in the
Arctic
18–20
, with vortex-average amounts at some altitudes similar to
those in the Antarctic and dramatically larger than the near-zerovalues
at that time in most Arctic winters. Furthermore, PSCs in 2011
spanned an altitude range comparable to that in the Antarctic, an
uncommon occurrence in the Arctic
18–20
. Particles in long-lasting
PSCs can grow large enough to sediment, resulting in denitrification,
permanent removal of HNO
3
from the stratosphere
7,12
. By late March
2011 no PSCs remained (Fig. 3a), yet HNO
3
mixing ratios were much
lower than observed in any previous Arctic winter (Fig. 2a). The con-
tinuing depression in HNO
3
after PSCs had evaporated indicates
denitrification. Albeit less severe than in typical Antarctic winters
(Fig. 2b, c, 3c), the extent and degree of denitrification in 2011 were
unmatched in the Arctic, approaching the range of Antarctic condi-
tions for the first time.
Decreasing HCl and increasing ClO signify chlorine activation
(Fig. 2d–i). Some ClO enhancement has occurred in all recent
Arctic winters, but has never been as prolonged and extensive as that
in 2011. In late February, high ClO pervaded the sunlit portion of the
vortex. The 2011 values vastly exceed the range previously observed in
the Arctic from late February through to the end of March. They also
briefly lie outside the Antarctic seasonal envelope, primarily because
the higher solar zenith angles of the Antarctic measurements used
here lead to ,30% lower ClO under fully activated conditions. In late
February, HCl values (unaffected by solar zenith angle issues) fall
along the lower boundary of the Antarctic envelope, confirming the
picture seen in ClO. The vertical extent of chlorine activation was also
comparable to that in the Antarctic (Fig. 3d, e).
In previous cold Arctic winters, chlorine was deactivated (converted
from ozone-destroying forms into less reactive reservoir species) by
mid-March
11
; even in 1997, ClO started to decline by late February
(Fig. 2g). In 2011, by contrast, ClO began decreasing rapidly only about
a week earlier than is typical in the Antarctic. ClO data in late February
1997 indicate that not only were maximum values lower than those in
early March 2011, but also the vertical range of enhancement was
shallower, with weaker activation at low altitudes than in 2011
(Fig. 3e), consistent with the higher altitudes and decreasing extent
(Figs 1b, 3b, Supplementary Fig. 2) of T , T
act
.
When chlorine is deactivated, whether it is converted first into HCl
or ClONO
2
depends sensitively upon HNO
3
and ozone abundances. In
the Arctic, chlorine is normally deactivated through initial reformation
of ClONO
2
. In the severely denitrified and ozone-depleted Antarctic
vortex, production of ClONO
2
is suppressed and that of HCl highly
favoured
11,12,21
. In March 2011, the recovery of HCl followed a much
more Antarctic-like pathway than has been observed in any other
Arctic winter.
The largest Arctic chemical ozone loss was previously observed in
2005, followed closely by 2000 and 1996
7,12–14
. Although low tempera-
tures persisted until the end of March 1997, the ozone loss in that year
was far less. No previous year rivals 2011, when the evolution of Arctic
ozone more closely followed that typicalof the Antarctic (Fig. 2j). Ozone
profiles in late March 2011 resemble typical Antarctic late-winter pro-
files much more strongly than they do the average Arctic one (Fig. 3f).
Because mixing in April 2011 (for example, lamination events larger
than that shown in Fig. 3f) entrained ozone-rich air into the vortex, the
slight decrease in vortex-averaged ozone at a potential temperature of
485 K from 26 March to 20 April (from ,1.8 to ,1.6 p.p.m.v., Fig. 2j)
indicates continuing chemical loss during this interval.
Estimates of chemical ozone loss
Chemical loss is difficult to quantify in the Arctic, where transport
from above replenishes ozone in the lower stratospheric vortex,
obscuring the signature of chlorine-catalysed destruction
12,22,23
. The
evolution of the long-lived trace gas nitrous oxide (N
2
O) reflects steady
downward transport throughout the 2010–11 winter–spring, indi-
cating that subsidence partially masked chemical loss. Horizontal
transport can also confound the signature of chemical loss, bringing
air into the vortex that has either higher
24
or lower
14
concentrations of
ozone, depending on the altitude and latitude from which it originates.
Representative results from two types of chemical loss calcula-
tions
24–28
based on balloon-borne and satellite observations are shown
in Fig. 4. The differences (up to ,0.4 p.p.m.v. at the end of March
2011) in estimates derived from the various methods and data sets
imply some uncertainty in the chemical loss determination. Year-to-
year differences in the amount of ozone loss are very similar when
obtained from any method/data set combination, however, indicating
a high degree of precision in the relative amount of calculated loss
0.0
0.1
0.2
0.3
0.4
Max. PV gradient (10
–4
(s deg)
–1
)
1 Jun. 1 Jul. 1 Aug. 1 Sep. 1 Oct.
a PV gradient, 1979–2011
0.0
0.1
0.2
0.3
0.4
0.5
0.6
V(T<T
act
)/V(vortex)V
psc
/V(vortex)
1 Dec. 1 Jan. 1 Feb. 1 Mar. 1 Apr.
1980 1985 1990 1995 2000 2005 2010
0.0
0.1
0.2
0.3
b Vortex fraction of low temperature air, 1979–2011
c Winter mean vortex fraction of low temperature air
Year
Figure 1
|
Meteorology of the Arctic lower stratosphere. a, Vortex strength
(as indicated by maximum potential vorticity
49
(PV) gradients) at 460 K
potential temperature (,18 km altitude, ,65 hPa level). b, Fraction of vortex
volume at potential temperatures between 390 and 550 K with a temperature
less than the chlorine activation threshold (T
act
). Light (dark) grey shading
shows range of Arctic (Antarctic) values for 1979–2010. Antarctic dates are
shifted by six months (top axis in a) to show the equivalent season. c,Winter
mean V
psc
during the past 32 years, expressed as a fraction of vortex volume.
Red, orange, green, purple and blue lines/bars show the 2010–11, 2004–05,
1999–2000, 1996–97 and 1995–96 Arctic winters, respectively.
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between different years. Chemical destruction was severe between
,16 and 22 km altitude, with the largest loss exceeding 2.5 p.p.m.v.
by 26 March 2011 (Fig. 4a). By 31 March 2011, chemical loss was
nearly double that in 2005 from ,18 to above 22 km, and similar to
that in 2005 at lower altitudes (Fig. 4b, c). From ,18 to 20 km, more
than 80% of the ozone present in January had been chemically
destroyed by late March. Chemical removal in 1996 and 2000 started
at a rate similar to that in 2011 (Fig. 4c), but ceased by late March;
maximum losses in 2000 approached those in 2011, but extended over
a much smaller vertical range (Fig. 4b). Loss in 1996, 2000 and 2005
considerably exceeded that in 1997, with greater destruction at lower
altitudes in those years contributing more to total column loss
7,12,13
.
Chemical loss in 2011 was two to three times larger than that in 1997,
and about twice that in 1996 and 2005 above ,16 km; from ,15 to
23 km it was comparable to that in the Antarctic ozone hole in 1985
29
.
Single ozone-sonde station measurements in early April 2011 suggest
continuing ozone loss (Fig. 4c).
Although the meteorology during March–April was similar in 1997
and 2011, ozone loss was much more pronounced in 2011. Photo-
chemical box model simulations (Supplementary Fig. 4, Supplemen-
tary Discussion) elucidate how early winter conditions set the stage for
record springtime ozone destruction in 2011. Chlorine activation
brought on by enduring cold from December through to the end of
February led to ,0.7–0.8 p.p.m.v. lower ozone at the beginning of
March 2011 (Figs 2j, 4c). The early onset of continuous cold also
facilitated formation of PSC particles large enough to sediment, result-
ing in ,4 p.p.b.v. less HNO
3
by March in 2011 than in 1997 (Fig. 2a).
The degree of denitrification has a profound impact on the severity of
springtime Arctic ozone loss
30
. By delaying chlorine deactivation,
lower HNO
3
by 1 March was responsible for ,0.6 p.p.m.v. more ozone
2
4
6
8
10
12
14
Vortex average HNO
3
(p.p.b.v.)
1 Jun. 1 Jul. 1 Aug. 1 Sep. 1 Oct.
a HNO
3
d HCl
g ClO
j O
3
b
26 Mar.
Arctic
c
Antarctic
3.0
4.0
5.0
6.0
7.0
0.0
0.5
1.0
1.5
2.0
2.5
Vortex average HCl (p.p.b.v.)
e
28 Feb.
f
29 Aug.
28 Feb. 29 Aug.
0.24
0.48
0.72
0.96
1.20
0.0
0.2
0.4
0.6
0.8
1.0
1.2
Vortex average ClO (p.p.b.v.)
h
i
0.30
0.60
0.90
1.20
1.50
1.0
1.5
2.0
2.5
3.0
3.5
Vortex average O
3
(p.p.m.v.)
HNO
3
(p.p.b.v.) HCl (p.p.b.v.) ClO (p.p.b.v.) O
3
(p.p.m.v.)
1 Dec. 1 Jan. 1 Feb. 1 Mar. 1 Apr.
k
l
1.2
1.4
1.6
1.8
2.0
2.2
2.4
2.6
26 Sep.
26 Mar. 26 Sep.
Figure 2
|
Chemical composition in the lowerstratosphere. a–l, Maps (right)
and vortex-averaged time series (left) at 485 K potential temperature (,20 km,
,50 hPa) for four different gases: HNO
3
(a, b, c), HCl (d, e, f), ClO (g, h, i) and
O
3
(ozone; j, k, l); mixing ratios from Aura MLS are shown. Averaging for the
time series is done within the white contour shown on the maps. Blue (purple)
triangles on time series, 1995–96 (1996–97) values from UARS MLS. Line
colours/shading as in Fig. 1, but shading is for Aura MLS measurements from
2005–10. Antarctic dates are shifted by six months (top axis on time series) to
show the equivalent season. Vertical lines show dates of maps in 2011 (2010) in
the Arctic (Antarctic). Black overlays on HNO
3
maps, T
act
(,196 K at this
level); HNO
3
may be sequestered in PSCs at lower temperatures. Dotted black/
white contour on ClO maps, 92u SZA, poleward of which measurements were
taken in darkness. Yellow/black triangles on ozone maps, locations of the
profiles in Fig. 3.
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loss after that date in 2011 than in 1997 (Supplementary Fig. 4,
Supplementary Discussion). The effects of denitrification and early-
winter loss together account for the disparity in ozone depletion in
these two winters (,1.5 p.p.m.v. more loss at 460 K in 2011 than in
1997, Fig. 4c, Supplementary Fig. 4). Loss as severe as that in 2011 thus
requires T , T
act
, with consequent chlorine activation and ozone
destruction, early in winter (as in 1996, 2000 and 2005, but not in
1997), a cold period and region before March sufficient to allow wide-
spread denitrification, and the persistence of a cold polar vortex into
April (as in 1997, but not in 1996, 2000 or 2005).
0 2 4 6
15
20
24
29
Altitude (km)
PSCs/aerosols
(10
–5
km
–1
sr
–1
)
a
190 210
b
Temperature (K)
0 2 4 6 8
HNO
3
(p.p.b.v.)
c
0.0 1.0 2.0
HCl (p.p.b.v.)
d
0.0 0.8 1.6
ClO (p.p.b.v.)
e
0 2 4 6
100
10
Pressure (hPa)
O
3
(p.p.m.v.)
f
Figure 3
|
Vertical composition information. a, Red, PSCs/aerosol amounts
averaged in the vortex over a week centred around 25 February 2011; dark blue,
the average for the same week in 2007–10; grey, the average over the equivalent
period (centred on 28 August) for the Antarctic in 2006–10; lavender, the Arctic
average for a week centred around 26 March 2011. (In late winter–spring,
maximum PSC altitudes are generally higher in the Arctic because early winter
PSC activity redistributes HNO
3
and water vapour to lower altitudes in the
Antarctic
18
). b–f, Daily average profiles of MERRA temperatures (b) and MLS
HNO
3
(c), HCl (d), ClO (e) and ozone (f). Red lines, data from a 4u 3 15u
latitude3 longitude box around 79u N, 12u E; in c, f, taken on 26 March; in
b, d, e, on 6 March 2011. Lavender, 7-day average for 2005–10 (1980–2010 for
b) centred on the same location and days. Grey, profiles in a similar box in the
Antarctic (79u S, 12u E) on 26 September for c, f, and on 8 September 2010 for
b, d, e. Dotted black line in b, approximate T
act
(195 K), see text. Purple line in
b, 7-day average around 6 March 1997, centred on same location. Purple line in
e, a midday ClO profile from UARS MLS on 26 February 1997 averaged in an
8u 3 30u box centred at the same Arctic location. A high-resolution ozone-
sonde profile at Ny A
˚
lesund on 26 March 2011 (black in f) agrees well with
MLS; lamination, a signature of mixing with ozone-rich extra-vortex air, is
apparent as a local maximum near 60 hPa.
380
420
460
500
540
Potential temperature (K)
1 Jan. 1 Feb. 1 Mar. 1 Apr.
a
0 1 2 3 4
Chemical ozone loss
(p.p.m.v.)
15
17
19
21
Approximate altitude (km)
b
1 Jan. 1 Feb. 1 Mar. 1 A
p
r.
0
1
2
3
4
Ozone (p.p.m.v.)
NH 2010–11
NH 2004–05
NH 1999–2000
NH 1998–99
NH 1996–97
NH 1995–96
SH 2003
SH envelope
c
0.5
1.0
1.5
2.0
Ozone loss (p.p.m.v.)
Figure 4
|
Chemical ozone loss estimates. a, Chemical loss as a function of
time and potential temperature from passive subtraction of MLS and ATLAS
passively-transported ozone (initialized with December MLS data).
b, Chemical loss from ozone sondes in unmixed vortex air as a function of
‘spring equivalent potential temperature’
48
(black contours in a). Shading,
Antarctic range defined by 1985 (the first year with profile measurements
inside the ozone hole
29
) and 2003 (a recent year with a severe ozone hole). The
2003 Antarctic curve is shifted by six months minus 10 days because ozone
sondes that year predominantly sampled the outermost vortex, where ozone
loss begins earliest. c, Ozone at a spring equivalent potential temperature of
465K (white contour in a), near the level of maximum chemical loss. Shading,
the region below the minimum reached in the 1985 Antarctic ozone hole. In
April 2011 most soundings sampled the disturbed vortex edge; only two were
made in air uninfluenced by mixing (red dots). Error bars, 1s uncertainties
based on the scatter of individual ozone-sonde measurements. Line colours as
in Fig. 1; 1998–99 (a winter with no ozone loss) is shown in cyan. NH, Northern
Hemisphere; SH, Southern Hemisphere.
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Column ozone
Total column ozone is a predominant factor determining exposure of
Earth’s surface to ultraviolet radiation
7,12
. In the context of previous
Arctic winters, 2011 was truly remarkable: the fraction of the Arctic
vortex in March with total ozone less than 275 Dobson units (DU) is
typically near zero, but reached nearly 45% in 2011 (Fig. 5a). Because
of the dynamically-driven correlation between total ozone and lower-
stratospheric temperature
23,31–34
(Supplementary Discussion), the
abiding cold in 1997 and 2011 would have led to lower March total
ozone than in other Arctic winters even without chemical loss;
dynamical conditions in March–April 1997 particularly favoured
low total ozone
33
(Supplementary Discussion). In March 2011, however,
the area of low total ozone covered more than twice as much of the
vortex as in 1997, and the daily vortex ‘ozone deficit’ (Supplementary
Fig. 5a) was 30–50 DU larger, consistent with the greater chemical loss
(Fig. 4). Maximum 2011 vortex fractions of low ozone approached
those in early Antarctic ozone holes (Fig. 5a). The close correspond-
ence between the vortex and both low total ozone and the large Arctic
total ozone deficit (Fig. 5b, d) implies that low total ozone in March
2011 resulted primarily from chemical loss
31,32
(Supplementary
Discussion). The ozone deficit in the Antarctic (Fig. 5e) shows a
maximum over 0–90u W, and a minimum over 90–200u E, reflecting
a vortex position in 2010 different to that in the reference state (which
is less robust than that for the Arctic). Differences in morphology deep
in the vortex are, however, minimal. The 2011 Arctic ozone deficit was
at least comparable to that in the 2010 Antarctic vortex core at an
equivalent time.
An echo of the Antarctic
In the absence of chemical ozone loss, downward transport during
winter results in a springtime maximum in total ozone; because this
transport is stronger in the Arctic, background ozone levels there are
,100 DU higher than those in the Antarctic
7,23
. Therefore Arctic
spring total ozone could, even after chemical destruction comparable
to that in an Antarctic ozone hole (commonly defined by values less
than 220 DU; refs 7, 12), exhibit only a weak maximum in total ozone
rather than a well-defined minimum. Examination of the long-term
ozone-sonde record in the Arctic shows that abundances near 250 DU
or less are well below typical autumn values, thus appearing as a ‘hole’
in total ozone. Dynamical processes can result in transient regions of
very low total ozone (Supplementary Discussion, Supplementary Figs
5, 6) and/or local minima in lower-stratospheric ozone profiles (for
example, via ozone-poor extra-vortex air transported into the polar
vortex
14,24
). For an interhemispheric comparison of chemical loss, it is
thus important to verify that observed Arctic ozone decreases were
primarily related to chemical, rather than dynamical, processes.
Figure 4 shows that the precipitous decline in Arctic ozone in
February–March 2011 resulted from chemical loss of similar mag-
nitude to that in the Antarctic in the mid-1980s. Observed ozone
between ,15 and 20 km altitude decreased to values matching the
minima in early Antarctic ozone holes and those reached at the cor-
responding time in some recent Antarctic winters (Figs 2j–l; 3f). In
late March–early April, most ozone-sonde profiles in the vortex had
mixing ratios less than 1 p.p.m.v., with values ,0.7 p.p.m.v. over an
approximately 2-km altitude region, and some dipping to 0.5 p.p.m.v.
(Supplementary Fig. 7). Minimum total ozone in spring 2011 was
continuously below 250 DU for ,27 days (Supplementary Fig. 5b),
with a maximal area below that level of ,2 3 10
6
km
2
(roughly five
times the area of Germany or California). Values dropped to ,220–
230 DU for about a week in late March 2011.
In these respects, chemical ozone destruction in the 2011 Arctic
polar vortex attained, for the first time, a level clearly identifiable as an
Arctic ozone hole. On the other hand, although the magnitude of
chemical depletion was comparable to that in the Antarctic, total
ozone values remained higher and, because the areal extent of the
Arctic vortex was much smaller (,60% the size of a typical
Antarctic vortex), the low-ozone region was more confined.
The Arctic winter stratosphere exhibits striking interannual vari-
ability. The past decade has included the four most dynamically active
(hence among the warmest) Arctic winters in the past 32 years (ref.
35) and now the two coldest winters with largest ozone loss
7,12–14
,
extending the previously noted trend of the coldest winters becoming
colder
13,16
. Had implementation of the Montreal Protocol not curbed
the increase in stratospheric halogen loading, formation of an Arctic
ozone hole would have already become common even in moderately
cold winters
36
. Even with the lower anthropogenic halogen levels
actually reached, the potential for Antarctic-like ozone loss in the
Arctic in the event of a persistently cold winter–spring such as that
in 2010–11 has been recognized for decades
5,22
. Despite temperatures
that were generally far higher than those in Antarctic winter, Arctic
chemical ozone destruction in 2011 rivalled that in some Antarctic
ozone holes. The development of an Arctic ozone hole under condi-
tions only slightly more extreme than those in some previous Arctic
winters raises the possibility of yet more severe depletion as lower-
stratospheric temperatures decrease. More acute Arctic ozone
destruction could exacerbate biological risks from increased ultra-
violet radiation exposure, especially if the vortex shifted over densely
populated mid-latitudes, as it did in April 2011.
Our present understanding of what drives variability in the Arctic
winter stratosphere is incomplete. Stratospheric temperatures and
vortex evolution depend on the atmosphere’s radiative properties
and propagation of wave activity
37,38
, which are being modified by
increasing greenhouse gas concentrations. Day-to-day tropospheric
disturbances can lead to stratospheric warming or cooling, depending
0.0
0.1
0.2
0.3
0.4
0.5
0.6
0.7
Area(<275 DU) / Area(vortex)
1 Feb. 1 Mar. 1 Apr.
1 Aug. 1 Sep. 1 Oct.
a Vortex fraction of low total O
3
b
26 Mar.
c
26 Sep.
280
320
360
400
440
Column O
3
(DU)
d
26 Mar.
e
26 Sep.
–110
–70
–30
10
50
Column decit (DU)
Figure 5
|
Total column ozone. a, Time series of the fraction of 460 K vortex
area with total ozone below 275 Dobson units (DU) in February–April in the
Arctic (bottom axis), and in August–October in the Antarctic (top axis). Line
colours/shading as in Fig. 1. 2005–2011 values are from OMI; earlier values are
from TOMS (Total Ozone Mapping Spectrometer) instruments
50
. Maps show
OMI total ozone (b, c) and ozone deficit (d, e) in the Arctic (Antarctic) on
26 March 2011 (26 September 2010). Overlays as in Fig. 2 but at 460 K.
ARTICLE RESEARCH
00 MONTH 2011 | VOL 000 | NATURE | 5
Macmillan Publishers Limited. All rights reserved
©2011

on their geographical location and the stratospheric vortex structure,
which controls their upward propagation
39,40
. Current climate models
do not fully capture either the observed short-timescale patterns of
Arctic variability or the full extent of the observed longer-term cooling
trend in cold stratospheric winters; nor do they agree on future cir-
culation changes that affect trends in transport
41,42
. Our ability to
predict when conditions similar to, or more extreme than, those in
2011 may be realized is thus very limited. Improving our predictive
capabilities for Arctic ozone loss, especially while anthropogenic
halogen levels remain high, is one of the greatest challenges in polar
ozone research. Comprehensive stratospheric data sets, such as those
used here, are critical to meeting that challenge.
METHODS SUMMARY
MERRA (Modern Era Retrospective-analysis for Research and Applications
43
)
fields are used for temperature and vortex analysis and for vortex averaging of
composition measurements. The CALIOP (Cloud-Aerosol Lidar with
Orthogonal Polarization) on the CALIPSO (Cloud-Aerosol Lidar and Infrared
Pathfinder Satellite Observations) satellite
44
provides PSC/aerosol information.
Trace gas profiles are from the Microwave Limb Sounder (MLS)
45
on NASA’s
Aura satellite. Only daytime ClO measurements are used. Northern (southern)
high latitudes are sampled near midday (in late afternoon), thus the average solar
zenith angle (SZA) of MLS Antarctic measurements is ,7u higher than that in the
Arctic. Reactive chlorine partitioning shifts away from ClO at higher SZAs
7,12
,
leading to ,30% lower ClO measured in the Antarctic than in the Arctic under
fully activated conditions. An instrument anomaly disrupted MLS measurements
from 27 March to 20 April 2011. UARS (Upper Atmosphere Research Satellite)
MLS measurements, used for 1995–1996 and 1996–1997 analyses, are sparse
because of the UARS yaw cycle and other measurement gaps
26
.
Total column ozone is measured by the Dutch-Finnish Ozone Monitoring
Instrument (OMI)
46
on Aura. Total ozone ‘deficit’ is the difference between daily
values and a reference that is minimally affected by chemical loss.
Measurements from MLS and the Match network of balloon-borne ozone
soundings (ozone sondes)
47
are used to estimate chemical ozone loss in two ways.
The difference between calculated ‘passive’ (influenced only by transport) ozone
and observed ozone is computed, with passive ozone obtained using MLS nitrous
oxide
14
, a ‘reverse trajectory’ model
25,26
, and the ATLAS (Alfred Wegener Institute
Lagrangian Chemistry/Transport System) model
27
. Vortex ozone is also examined
on the surfaces on which it subsides
12,14,28,48
, with descent rates from modelled
radiative heating/cooling rates averaged over the polar vortex
48
.
Photochemical box model runs were performed using the chemical model
from ATLAS
27
to test the sensitivity of ozone loss to initial ozone amounts and
denitrification.
Full Methods and any associated references are available in the online version of
the paper at www.nature.com/nature.
Received 3 May; accepted 7 September 2011.
Published online 2 October 2011.
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Supplementary Information is linked to the online version of the paper at
www.nature.com/nature.
Acknowledgements We thank the MLS (especially A. Lambert, D. Miller, W. Read,
M. Schwartz, P. Stek, J. Waters), OMI (especially P. K. Bhartia, G. Jaross, G. Labow),
CALIPSO and Match science teams, as well as A. Douglass, J. Joiner and the Aura
project, for their support. We also thank W. Daffer and R. Fuller for programming
assistance at JPL; the many observers whose work went into obtaining the ozone-sonde
measurements; the ozone scientists who participated in the discussion of the 2011
Arctic ozone loss and appropriate definition of an Arctic ozone hole (including, but not
limited to, N. Harris, G. Bodeker, G. Braathen, M. Kurylo, R. Salawitch); and especially
P. Newman and K. Minschwaner for discussions and comments. Meteorological
analyses were provided by NASA’s Global Modeling and Assimilation Office (GMAO)
and by the European Centre for Medium-Range Weather Forecasts. We thank
S. Pawson of GMAO for advice on usage of the MERRA reanalysis. Ozone-sonde
measurements at Alert, Eureka, Resolute Bay, Churchill and Goose Bay were funded by
Environment Canada. Additional ozone sondes were flown at Eureka as part of the
Canadian Arctic Atmospheric Chemistry Experiment (ACE) Validation Campaign and
were funded by the Canadian Space Agency. Academy of Finland provided partial
funding for performing and processing ozone-sonde measurements in Jokioinen and
Sodankyla
¨
. Ozone soundings and work at AWI were partially funded by the EC DG
Research through the RECONCILE project. Work at the Jet Propulsion Laboratory,
California Institute of Technology, and at Science Systems and Applications Inc., was
done under contract with NASA.
Author Contributions G.L.M. and M.L.S. led analysis of MLS data; M.R. led analysis of
ozone-sonde data; G.L.M. led the meteorological data analysis. M.R., G.L.M., N.J.L. and
I.W. did chemical ozone loss calculations. R.L. and M.R. performed and analysed
chemical box model calculations. M.C.P. and L.R.P. provided CALIPSO/CALIOP data
analyses; E.R.N. and P.V. provided TOMS and OMI data analyses. L.F., M.L.S., G.L.M. and
N.J.L. provided expertise on MLS data usage; D.P.H., P.V. and P.F.L. provided expertise
on OMI data usage. J.D., V.D., H.G., B.J., R.K., E.K., N.L., A.M., C.T.M., H.N., M.C.P., D.W.T.,
P.v.d.G., K.A.W. and N.S.Z. were responsible for performing and processing
ozone-sonde measurements. All authors contributed comments on the manuscript.
G.L.M., M.L.S. and M.R. jointly compiled and synthesized the results. G.L.M. and M.L.S.
wrote the paper.
Author Information CALIOP data are publicly available at http://eosweb.larc.nasa.gov/
PRODOCS/calipso/table_calipso.html, MLS data at http://disc.sci.gsfc.nasa.gov/Aura/
data-holdings/MLS, OMI data at http://disc.sci.gsfc.nasa.gov/Aura/data-holdings/
OMI/omto3_v003.shtml, and GEOS-5 MERRA analyses through http://
disc.sci.gsfc.nasa.gov/mdisc/data-holdings/merra/. The balloon-borne Antarctic
ozone-sonde data recorded in 1985 and the following years are publicly available at
http://dx.doi.org/10.1594/PANGAEA.547983. Reprints and permissions information
is available at www.nature.com/reprints. The authors declare no competing financial
interests. Readers are welcome to comment on the online version of this article at
www.nature.com/nature. Correspondence and requests for materials should be
addressed to G.L.M. (Gloria.L.Manney@jpl.nasa.gov) or M.L.S.
(Michelle.L.Santee@jpl.nasa.gov).
ARTICLE RESEARCH
00 MONTH 2011 | VOL 000 | NATURE | 7
Macmillan Publishers Limited. All rights reserved
©2011

METHODS
Data sets. Modern Era Retrospective-analysis for Research and Applications
(MERRA)
43
fields, from the Goddard Earth Observing System Version 5.2.0
(GEOS-5) data assimilation system, are used for the temperature and vortex
analysis. The Cloud-Aerosol Lidar with Orthogonal Polarization (CALIOP) on
the Cloud-Aerosol Lidar and Infrared Pathfinder Satellite Observations
(CALIPSO) satellite
44
provides PSC/aerosol information. CALIOP measure-
ments began in April 2006. Trace gas profile measurements are from the
Microwave Limb Sounder (MLS)
45
on NASA’s Aura satellite, and the predecessor
MLS instrument
26
on the Upper Atmosphere Research Satellite (UARS). Total
column ozone data are from the Dutch-Finnish Ozone Monitoring Instrument
(OMI)
46
on board Aura. The historical total ozone record comprises data from
Nimbus-7 and Earth Probe Total Ozone Mapping Spectrometer (TOMS)
50
. Aura
MLS and OMI measurements are available from August 2004 through to the
present. UARS MLS measurements were obtained from September 1992 through
to early 2000, with increasingly sparse sampling in the later years
26
. TOMS data
are available beginning in 1979, but no TOMS instrument was taking measure-
ments during the 1995–96 Arctic winter.
Measurements from the Match network of balloon-borne ozone soundings
(ozone sondes)
47
are used in some of the chemical ozone loss estimates.
Temperature and vortex analysis. Potential vorticity
49
(PV) is used to define
the vortex, with a contour of ‘scaled’ PV of 1.4 3 10
–4
s
21
(in vorticity units)
demarking the vortex edge
51,52
. Vortex strength is diagnosed as the maximum
daily gradient in PV as a function of equivalent latitude (the latitude that would
enclose the same area between it and the pole as a given PV contour)
51–53
. Scaled
PV multiplied by 10
4
is used in the calculation, resulting in units for its gradient of
10
24
(s degrees equivalent latitude)
21
.
The temperature threshold for chlorine activation, T
act
, is estimated using the
formula for nitric acid trihydrate formation
54
, which depends on pressure, HNO
3
and H
2
O. Climatological HNO
3
and H
2
O profiles are used, derived from UARS
data. The area with T , T
act
is calculated on seven isentropic surfaces in the lower
stratosphere: 390, 410, 430, 460, 490, 520 and 550 K; T
act
on these levels is 197.5,
197.2, 196.8, 196.5, 195.9, 195.3 and 194.5 K, respectively. To get the volume with
T , T
act
from 380 through 565 K, the areas at each of the seven levels are mul-
tiplied by the estimated altitude associated with that layer and summed. The
altitude range associated with each layer is obtained from a standard potential
temperature profile as a function of altitude derived from high latitude temper-
ature soundings taken during the 1988–89 through to 2001–02 winters (the same
profile was used for V
psc
calculations in refs 13, 16 and 48). These thicknesses are
1.29088, 1.19995, 1.36770, 1.46281, 1.30554, 1.18199 and 1.07382 km for the
seven levels listed above. Vortex volume is calculated from vortex area in the
same manner. Winter mean V
psc
is calculated over 16 December through to 15
April. Previous studies have shown that V
psc
scaled by the vortex area is a good
proxy for chlorine activation and ozone loss potential
17
. Additional temperature
and vortex diagnostics are described in Supplementary Information.
Polar stratospheric cloud and aerosol information. Particulate backscatter
averaged over the polar vortex derived from CALIOP data is used to provide
PSC/aerosol information. Total attenuated backscatter at 532 nm, b(z), is one of
the basic CALIOP Level 1B data products. b(z) is the sum of the particulate
backscatter (due to liquid aerosol and PSCs), b
p
(z), and molecular backscatter,
b
m
(z). b
m
(z) is calculated using GEOS-5 molecular density profiles (included in
the CALIOP Level 1B data files) and a theoretical value for the molecular scatter-
ing cross-section
55
. Profiles of b
p
(z) are then produced by subtracting b
m
(z) from
b(z). Vortex-averaged profiles of b
p
(z) are produced by averaging all CALIOP
b
p
(z) profiles located inside the vortex edge (defined using information available
in GEOS-5 Derived Meteorological Product (DMP) files for the nearly-coincident
Aura MLS data
52
) over the selected time interval.
MLS trace gas profile measurements and analysis. Trace gas profile measure-
ments of HNO
3
, HCl, ClO, ozone and N
2
O (a long-lived tracer used to assess
descent) are from Aura MLS
45
version 3 retrievals; data quality screening is as
recommended in the MLS data quality document
56
. MLS data are retrieved on
pressure surfaces; potential temperature as a function of pressure from MLS
DMPs
52
calculated from GEOS-5 analyses is used to interpolate to isentropic
surfaces. Vortex averages of MLS data are calculated using the 1.4 3 10
24
s
21
scaled PV contour to define the vortex edge, using PV values from the MLS
DMPs
52
. Active chlorine is in the form of ClO mainly during the daytime, and
thus measured ClO amounts vary with the solar zenith angle (SZA) at which the
measurements are taken. Only daytime ClO measurements are used here.
Northern high latitudes are sampled near midday local time, southern high
latitudes are sampled in late afternoon, thus the SZA of Aura MLS Antarctic
measurements is ,7u higher on average than that in the Arctic. Reactive chlorine
partitioning shifts away from ClO at higher SZAs
7,12
, leading to ,30% lower ClO
measured by Aura MLS in the Antarctic than in the Arctic under fully activated
conditions. MLS measurements are unavailable from 27 March through to
20 April 2011 because of an instrument anomaly. Upper Atmosphere Research
Satellite (UARS) MLS measurements, used for analysis of 1995–96 and 1996–97,
are sparse because of the UARS yaw cycle and other measurement gaps
26
. The
time of day of UARS measurements varied through the yaw cycle, in the middle of
which no daytime ClO measurements were obtained
10
; thus ClO values shown in
1995–96 and 1996–97 near those dates (including the mid-February 1996 mea-
surements shown in Fig. 2g) are not representative of the degree of chlorine
activation.
Chemical loss calculations. Chemical ozone loss is quantified by two methods,
both widely used for such calculations
7,12,24–28,47,48
. In the ‘passive subtraction’
method
25–27
, a transport model is used to calculate the evolution of ozone in
the absence of chemical changes (‘passive’ ozone). The difference between passive
ozone and observed ozone provides an estimate of chemical loss.
Here, passive ozone is obtained in three different ways. First, MLS observations
of N
2
O, a long-lived species unaffected by chemical processes, are used to cal-
culate vertical motion, and that estimate of descent is then used to calculate how
initial MLS ozone profiles would have evolved in the absence of chemical loss
14
.
Second, a ‘reverse trajectory’ transport model
25,26
is used to transport an initial
state based on MLS-observed ozone with no chemistry. Finally, the ATLAS
(Alfred Wegener Institute Lagrangian Chemistry/Transport System) chemistry
and transport model is run in passive mode
28
, initialized with MLS ozone.
Vortex ozone is also examined in relation to the surfaces on which it is sub-
siding
12,14,28,48
. The descent rates used here are obtained by averaging radiative
heating/cooling rates from the radiation calculation used in the ATLAS model
over the polar vortex
48
. These rates are then used to examine vortex-averaged
MLS and ozone-sonde data on surfaces of ‘spring equivalent potential temper-
ature’
48
, defined as the potential temperature at which air originating at a given
level arrived at the end of March. Since the air descended on these surfaces, ozone
would have been constant on each such surface in the absence of chemical loss.
The ozone-sonde data used here are all from electrochemical concentration cell
(ECC) sondes, made by different manufacturers. Ozone-sonde data quality was
assessed in an intercomparison experiment
57
and is discussed in ref. 47. For
chemical loss calculations using ozone-sonde data, the profiles are first examined
using a procedure for detecting lamination in the profiles; such lamination (an
example is shown in Fig. 3f) is associated with mixing in of extra-vortex air, which
may obscure the signature of chemical loss. Profiles that have been significantly
altered by mixing processes, as indicated by lamination, are excluded from the
vortex averages used in the chemical loss calculations. 2010–11 Arctic ozone-
sonde data are provided as Supplementary Information.
Results from the ATLAS model passive subtraction calculations, and from the
calculations on spring equivalent potential temperature surfaces using the Match
network ozone-sonde data, are shown in Fig. 4; all panels show vortex averages.
These results have been compared with the results from the other methods
described above. While absolute ozone values obtained from different methods/
data sets vary significantly (up to ,0.4 p.p.m.v. at the end of March 2011), the
year-to-year variations in chemical loss calculated using all three methods agree
closely, indicating a high degree of precision in the relative amount of calculated
loss between different years.
The Alfred Wegener Institute chemical box model, also used as the chemical
module in ATLAS, simulates 175 reactions between 48 chemical species in the
stratosphere
27,58
This model was used to perform conceptual runs (Supplemen-
tary Fig. 4), started on 1 March with identical initial mixing ratios of all species
except HNO
3
and O
3
. For these two species values corresponding to 1997
(3 p.p.m.v. O
3
, 10 p.p.b.v. HNO
3
) and 2011 (2.2 p.p.m.v. O
3
, 6 p.p.b.v. HNO
3
)
(compare Figs 2a and 4c) were combined to yield four sets of initial conditions.
Initial ClO
x
was 2 p.p.b.v., corresponding to the vortex-averaged ClO
x
derived by
ATLAS from MLS ClO measurements on 1 March 2011. An air parcel at 70u N,
460 K potential temperature, with a temperature of 193 K throughout March, was
used. Heterogeneous reactions took place on liquid aerosols, rather than solid
(nitric acid trihydrate, NAT) PSCs, since the widespread existence of the latter is
inconsistent with MLS observations of gas-phase HNO
3
values (Fig. 2a) larger
than those the microphysical module predicts if NAT is present. A sensitivity
run showed that sporadically occurring solid PSCs did not change the results
significantly.
Column ozone and ozone deficit calculation. OMI total ozone data were
processed with version 8.5 of the TOMS algorithm and have been extensively
validated
59
. TOMS data were processed with version 8 of the algorithm. The OMI
and TOMS total ozone data used in this study were averaged on a fixed global
1u 3 1u latitude 3 longitude grid. Averages were computed by area-weighting
observations based on the overlap of their instantaneous field-of-view with each
grid cell. Only data that satisfy quality criteria based on measurement path length
and algorithm diagnostic criteria were included in the averaged samples.
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Individual total ozone retrievals included in the samples are expected to have a
root-mean-squared error of 1–2%.
Total ozone ‘deficit’ is calculated as the difference between daily values and a
reference that is minimally affected by chemical ozone loss. The reference for the
Arctic is the daily mean over all Arctic winters from 1978–79 through to 2009–10,
from OMI starting in 2004–05 and from TOMS for earlier years
50
. The Antarctic
reference state is the daily mean of TOMS measurements for 1979 through to
1981. Because the Antarctic reference state is based on only three years’ data for
each day, variations in vortex position are not effectively averaged out; this
reference is thus less robust than that for the Arctic, so patterns in daily maps
may partially reflect differences in vortex position between the reference and the
focus day.
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