Introduction
There has recently been great interest in the variability of middle
atmospheric trace constituents at high latitudes in the late winter and early
spring. This interest has been fueled, in part, by the occurrence of
prolonged sudden stratospheric warmings (SSWs) which can perturb the
composition and structure of the stratosphere and mesosphere for many weeks
(Manney et al., 2008a, b, 2009). These so-called extended SSWs are
characterized by elevated stratopauses which reform near and above 80 km
(Siskind et al., 2007; Manney et al., 2009). During the recovery phase of
these extended events, the anomalous zonal wind flow alters the gravity wave
propagation to the mesosphere, thus perturbing the mean meridional
circulation and driving a dramatic descent of mesospheric air down to the
stratosphere (Siskind et al., 2010; Chandran et al., 2013). For example,
have shown that mesospheric air enhanced in nitric oxide (NO)
and depleted in water vapor (H2O) and methane (CH4) can descend from
near 90 km in early February down to 40 km by early April. Bailey et
al. (2014) focused on the 2013 SSW; other analogous events occurred in 2004,
2006 and 2009 (Manney et al., 2005, 2009; Randall et al., 2009). An
additional motivation for most of the above studies is the interest in
quantifying the extent to which the enhanced nitric oxide can cause
reductions in polar upper stratospheric ozone (Funke et al., 2014).
There has been less attention paid to what happens to these dramatic
perturbations as the spring progresses and the wintertime circulation
transitions into a summer pattern. It has long been recognized that the
winter to spring transition is characterized by a decay and breakdown of the
wintertime westerly jet and its eventual replacement by a zonal mean
easterly flow around the polar region. This is known as the stratospheric
final warming (SFW) (Hu et al., 2014). It has been observed that certain
remnants of wintertime dynamical (Hess, 1990) or chemical tracer features
(Orsolini, 2001; Lahoz et al., 2007) can persist well into the summer season.
Most recently, work has focused upon specific events, whereby the SFW can
occur rather abruptly with a significant late season planetary wave event
(Allen al., 2011; Siskind et al., 2015a; Fiedler et al., 2014). These
planetary waves can transport low-latitude anticyclonic air poleward. This
air can displace the winter polar vortex and then remain “frozen in” for a
period of weeks or longer in late spring and early summer .
Alternatively, this transition can occur gradually without significant wave
activity. In the former case, the upper mesosphere often experiences cooler
and wetter conditions which can lead to the early onset of the polar
mesospheric cloud season. In the latter case, the upper mesosphere
remains warmer and drier. Siskind et al. (2015a) showed that 2011 and 2013
were years with an abrupt winter to spring transition and 2008 was a spring
with negligible planetary wave activity. They used these years to define the
extremes in springtime planetary wave activity and associated temperatures.
From the above, we can define four general scenarios for the transition from
winter to summer based upon the combination of the two perturbations outlined
above. We can have a year with extended descent of mesospheric air (typically
the result of a extended SSW) or a winter with weak descent. These winters
can be followed by springs with either an abrupt planetary wave transition to
a summer circulation or with a slower gradual transition. The purpose of this
paper is to categorize the four possible combinations of these springtime
scenarios and how they are manifested in the variability of trace
constituents such as CH4, chlorine monoxide (ClO) and ozone (O3). Among
our results, we will show that under certain circumstances, the zonal mean
distribution of these trace constituents can be perturbed for many months
even into the autumn. This is important because while the summer upper
stratosphere is generally understood to be under radiative and photochemical
control (Andrews et al., 1987), we will show how the zonal mean composition
can be sensitive to dynamical changes that might have occurred over half a
year prior.
Observations and model
SOFIE and MLS data
Overview of upper stratospheric and lower mesospheric zonal mean
CH4 observed by SOFIE for the indicated years. SOFIE observes at only one
latitude per day in each hemisphere. This latitude has some variation from
year to year, but is typically near 82∘ at the equinoxes and near
65–66∘ at the solstices. The horizontal axis label, doy, is day of year.
Our primary data come from the Solar Occultation for Ice
Experiment (SOFIE) on the Aeronomy of Ice in the Mesosphere (AIM) satellite (Russell III et al., 2009)
and the Microwave Limb Sounder (MLS) (Santee et al., 2008; Froidevaux et al., 2008) on the Aura satellite .
SOFIE measures profiles of
temperature, aerosols (ice and meteoric smoke) and O3, H2O, CO2, CH4 and NO using
the solar occultation technique. Since the AIM satellite is in a sun-synchronous polar orbit, the latitude of the occultations approximately
tracks the terminator and is above 82∘ near equinox and near 65∘ at solstices. SOFIE acquires approximately 15 samples day-1, uniformly spaced
in longitude. The vertical resolution is about 2 km. Gordley et al. (2009) quote
a precision for the CH4 data of 10 ppbv at 70 km.
This work uses version 1.3 SOFIE data. SOFIE CH4 data have previously been presented by
and and were shown to vary in a manner consistent with the other tracers of mesospheric descent measured by SOFIE;
ongoing validation studies (Rong et al., 2016)
with the Atmospheric Chemistry Experiment suggest general agreement to within approximately 12 %. Here we emphasize the
relative year to year variations.
Like AIM, the Aura satellite is also in a sun-synchronous orbit. However,
unlike SOFIE, because MLS observes ClO and O3 in emission, data are
obtained over all latitudes up to about 82∘ N. We used Version 4.2
data. The MLS ozone was validated by and used in a study
of lower mesospheric photochemistry by . The ClO data have
been validated by and compared with ground-based data by
. Santee et al. (2008) show that the precision of the MLS
ClO decreases for pressures less than 2 hPa; however, since we only show
monthly averages, this is not a problem for the present study. It is also
common practice to subtract the nighttime data from the daytime data (Santee
et al., 2008; Nedoluha et al., 20011) in order to reduce systematic biases;
however, for the high-latitude spring/summer conditions shown here, there are
often no night periods. Thus a given monthly average was constructed using
data from all local times without any background subtraction. The vertical
resolution of the MLS ClO observation (3–4 km) is somewhat coarser than
SOFIE. We thus interpolated the SOFIE profile to the MLS grid.
The Whole Atmosphere Community Climate Model (WACCM)
We also compare some of our results with WACCM (Garcia et al., 2007). WACCM
is the high-altitude atmospheric component of the NCAR Community Earth System
Model version 1 (CESM1). In its standard configuration, WACCM has 66 vertical
levels from the ground to about 5.9 × 10-6 hPa (approximately
140 km geometric height) and a horizontal resolution of
1.9∘ latitude × 2.5∘ longitude. See Garcia et
al. (2007) for a detailed discussion of the model climatology and
parameterizations. This version of WACCM uses specified dynamics
provided by the Navy Operational Global Atmospheric Prediction System –
Advanced Level Physics High Altitude (NOGAPS-ALPHA) (Marsh, 2011; Sassi et
al., 2013). NOGAPS-ALPHA is the high-altitude extension of the then
operational Navy's weather forecast system up to about 90–92 km (Eckermann
et al., 2009). have already shown that the combination of
WACCM and NOGAPS-ALPHA (hereinafter called WACCM/NOGAPS) produced a
successful representation of the descent of enhanced upper mesospheric and
lower thermospheric nitric oxide (NO) and depleted CH4 into the upper
stratosphere/lower mesosphere. By contrast, WACCM nudged only up to 50–60 km
by the Modern Era Retrospective Analysis for Research and Applications
(MERRA) dataset did not (see also Randall et al., 2015). Since mesospheric
descent is so important for understanding our present results, we only use
WACCM/NOGAPS for this study. Unfortunately, of the 7 years considered
here (2008–2014), WACCM/NOGAPS is only available for the first 2. We thus
can not use it to reproduce all the variability seen in the SOFIE data.
However, by comparing summer results from 2009 with 2008, we can provide a
broader context to the latitudinal extent of the CH4 changes and their
effect on the chlorine and ozone chemistry of the upper stratosphere.
Results
Methane (CH4)
Comparison of time series of zonal mean CH4 mixing
ratio at 1.47 hPa. (a) SOFIE data for the indicated years. The data
have been grouped into 5-day bins. See Fig. 1 for a discussion of the
latitudes. (b) WACCM for 2008 (solid) and 2009 (dashed) at a single
fixed latitude of 75∘ N.
Our specific interest is to highlight the consequences of the variations in
upper stratospheric CH4 as observed by SOFIE and shown in Figs. 1 and 2.
These figures illustrate the great variability that occurs in CH4 each
winter and spring. Figure 1, which presents 6 years of SOFIE CH4, shows
that each year is characterized by the descent of low values of CH4 from
the mesosphere in the period from February to early April (roughly day 30 to
day 110). This descent is characterized by large interannual variability and
was strongest in 2009 and 2013. These were years with prolonged SSWs followed
by elevated stratopauses and have been covered in the literature (Manney et
al., 2009; Randall et al., 2009; Bailey et al., 2014). The difference between
2009 and 2013 is that in 2013, there was a large frozen-in anticyclone event
(FrIAC; Manney et al., 2006) that transported air with high values of CH4
to high latitudes (Siskind et al., 2015a), whereas in 2009, no such springtime disturbance was evident. This is clearly seen in Fig. 2 where the CH4
jumps from below 0.1 ppmv on day 100 to over 0.3 ppmv by day 120. Years
with a more moderate and shorter period of winter/early spring descent are
2010 and 2012. These 2 years did not have elevated stratopause events as in
2009 and 2013, but there were wintertime SSWs in both years, and Straub et
al. (2012) discussed the descent of dry air at high latitudes in the lower
mesosphere during the late winter of 2010. The springtime vortex breakdown
occurred relatively gradually over many weeks in March and April for both
2010 and 2012, and thus there was no transport of high CH4 in either
spring. These years ended up being close to 2009 in having low values of
CH4 persist into the summer. Even less mesospheric descent was seen in
2008 and the least descent was seen in 2011. The year 2011 was characterized by a
strong undisturbed stratospheric polar vortex (Manney et al., 2011). Then, in
early April (Day 95) of that year, the largest FrIAC of the 36-year MERRA
dataset was recorded (Allen et al., 2011; Thieblemont et al., 2013), causing
a significant jump in upper stratospheric CH4.
After the spring, there is a second period of decreasing CH4 in the summer
(most noticeable after day 200). This summertime decrease is due to
photochemistry (Funke et al., 2014), as the production of O(1D) and OH,
both of which oxidize CH4, peak at high summer latitudes in the upper
stratosphere (Letexier et al., 1988). Since the upper stratosphere at this
time of year is dynamically quiet, the year to year variability in summer
CH4 is driven by the winter- and springtime dynamics. This can be seen in
Fig. 2a, which compares time series of upper stratospheric CH4 for the
6 years shown in Fig. 1 plus 2014. The figure shows that the lowest summer
CH4 was generally in 2009; this is the direct consequence of the late
winter descent that persisted without interruption until early April. By
contrast, the highest summer CH4 was in 2011 which is the result of the
dynamically quiet winter followed by the FrIAC in early April that caused the
CH4 to almost double. The other 5 years are intermediate, although as
noted above, 2010 and 2012 are close to 2009. For all 7 years, once the
relative abundances of CH4 were established by 1 May (day 121), they
remained mostly unchanged with respect to each other until October (around
day 280). Figure 2b shows WACCM zonal mean CH4 results for 1.47 hPa at
the single latitude of 75∘ N for 2009 and 2008. The reason for
sampling WACCM at a single latitude is to test whether the slow seasonal
drift of the SOFIE occultation pattern from 65 to 82∘ might be
affecting our comparisons. While there are some differences in absolute
abundance between WACCM and SOFIE for the first 30–40 days when late winter
conditions still prevailed, after that, in spring and summer, the agreement
between WACCM at one latitude and SOFIE over a small range of latitudes is
excellent. Thus we can conclude that the latitude variation of the SOFIE
occultations can be neglected. This is not surprising since wave activity and
latitudinal gradients are relatively weak in summer.
Categorization of summer upper stratospheric CH4.
Category
Winter
Spring
CH4
Representative
descent
PW
value
year
1.
high
low
lowest
2009
2.
high
high
intermediate
2013
3.
low
low
intermediate
2008
4.
low
high
highest
2011
PW is planetary wave.
Table 1 presents an idealized categorization of how the summer level of
Arctic upper stratospheric CH4 can be placed in the context of the four
categories of wintertime descent and early spring dynamical variability. The
years 2008, 2009, 2011 and 2013 are most representative of these idealized
cases. The other years are more intermediate; as noted above, 2010 and 2012
were closer to 2009 in having relatively strong late winter descent of
mesospheric air and a relative absence of springtime wave activity (with its
associated horizontal transport of low-latitude air to polar latitudes; cf.
Siskind et al., 2015a). The year 2014 is closer to 2011. As seen in Fig. 2, there was
a 50 % increase in CH4 in late March 2014 and we have previously,
tentatively suggested that there was a FrIAC event in that spring (Siskind
et al., 2015a). Certainly this categorization is qualitative, not quantitative; however, we suggest that it provides a useful framework for analyzing the spring and
summer CH4 variability.
Chlorine monoxide (ClO)
Here we explore the chemical consequences of the CH4 variations
illustrated above. CH4 has long been known to play an important role in
partitioning stratospheric chlorine (Solomon and Garcia, 1984). Specifically,
the reaction Cl+CH4→HCl+CH3 means that active
chlorine (ClOx= Cl + ClO) should vary inversely with CH4. For example,
Siskind et al. (1998) documented an increase in upper stratospheric ClO
during the early years of the Upper Atmospheric Research Satellite (UARS)
mission which was explained as a direct consequence of the decrease in CH4
observed by Nedoluha et al. (1998). Froidevaux et al. (2000) observed a
general anticorrelation between variations in ClO and CH4 at 2 hPa in the
tropics. They showed that there should be an inverse relationship between ClO
and CH4.
Scatter plot of zonal mean, monthly averaged MLS ClO vs. SOFIE
CH4 at 1.47 hPa. The MLS data are sampled at the SOFIE occultation
latitude, the monthly averages of which are indicated in each panel. In the
upper right of each panel the linear correlation coefficients (r)
between each dataset for each month are given, as well as the slope of the linear fit (m) in
units of ppbv of ClO per ppmv of CH4.
Figure 3 shows that this anticorrelation also exists between high-latitude
CH4 and ClO at 1.47 hPa during the spring and summer. It plots monthly
averaged SOFIE CH4 against MLS ClO (sampled at the SOFIE occultation
latitudes) for the period May–August. Although there are only seven datapoints for each month (6 in May due to missing data in 2014), the linear correlation coefficients of
-0.92 to -0.97 are
highly statistically significant. Note there is a general increase in ClO from late spring to late summer. This is consistent with the seasonal
decrease in CH4 and was discussed by Considine et al. (1998). Concerning the year to year variability,
the highest summertime ClO for the 7-year period is in 2009.
This is a legacy of the strong uninterrupted descent which followed the January 2009 SSW. Other years with relatively high ClO include
2010 and 2012 which, as we have discussed, were also years similar to 2009 in their combination of winter descent and spring planetary waves.
The lowest summertime ClO is in 2011.
This is the result of the strong FrIAC event which occurred in April 2011. The general range of summer ClO which stems from the above winter/spring dynamical
variability is about 50 %.
Figure 3 also gives the slopes (m) of the linear fit between ClO and
CH4. It shows a tendency for progressively steeper (more negative) slopes
as the summer progresses and methane decreases. In general, all the values of
m are more negative than the value (-0.42 ppbv ppmv-1) quoted by
Nedoluha et al. (2011) for tropical conditions. However, Nedoluha et
al. (2011) make the point that the CH4 is relatively high in the tropics
(about 0.6 ppmv according to their Fig. 7). Thus the pattern of steeper
slopes for lower CH4 is robust across both Nedoluha et al's and our
analyses. This is precisely the pattern one would expect for the inverse
power relationships discussed by Froidevaux et al. (2000). Thus the present
SOFIE/MLS comparison is consistent with studies using both UARS and ground-based data that showed ClO and CH4 in the upper stratosphere varying with
a high degree of anticorrelation.
The color contours on the left are zonal mean WACCM/NOGAPS
difference fields for August 2009 minus August 2008 for ClO (top) and CH4
(bottom). The vertical dashed white line is the mean latitude of the SOFIE
occultations for August. On the right, a vertical profile of the model
difference at the SOFIE occultation latitude (solid line with plus symbols)
is compared with MLS ClO and SOFIE CH4 (data are dotted/dashed curves with
stars). Note that x axes for the right panels are reversed from one another
since the ClO change is positive, while the CH4 change is negative.
To get a broader picture of the ClO and CH4 changes at latitudes other
than the narrow range sampled by SOFIE, Fig. 4 shows the monthly average
zonal mean WACCM/NOGAPS ClO and CH4 difference fields for August 2009
minus August 2008. Profiles that are compared with MLS (for ClO) and SOFIE
(for CH4) for the SOFIE occultation latitude (given by the dashed white line
in the color panel) are also shown in the right-hand plots. The comparison
between the model and the data is excellent. Since the difference between
2009 and 2008 represents about half the difference between the extreme years
discussed above (2009 and 2011), one can multiply the ClO and CH4
difference values in Fig. 4 by a factor of 2 to get an estimate of the full
range. The model shows that the low 2009 CH4 and high 2009 ClO shown in
Fig. 4 are part of a broad region of perturbation extending from
40 to 50∘ N to the pole and covering the altitude region between about
1 and 8 hPa. There may be a small vertical offset, perhaps one grid point,
whereby the model profile is shifted slightly downward relative to both the
MLS and SOFIE data. A similar offset was recently noted by
in their WACCM/NOGAPS simulation of the 2009 descent of mesospheric NOx.
Since the summer CH4 depletion is a consequence of the winter descent,
this offset may reflect the small discrepancy seen by .
Figure 4 shows that the effect of the CH4 on ClO occurs over a relatively
deep layer in the upper stratosphere; the detailed plots of the time behavior
of CH4 and ClO, specifically Figs. 2 and 3, represent only the uppermost
edge of this larger perturbation. The reason for focusing on this narrower
region is that these altitudes, between 1 and 3 hPa, are where the chlorine
cycle is affecting the ozone. This is discussed in the next section.
Ozone (O3)
Time series of zonally averaged ozone from MLS at 75∘ N.
Figure 5 presents a time series of upper stratospheric ozone from MLS in a
format similar to Fig. 2 for CH4. Only 4 years are shown because in
summer, the curves almost overlap and it would be hard to distinguish all
7 years clearly. The 4 years shown correspond to the representative years
given in Table 1. The figure shows very large variability in March and April,
both intra- and interannually. This is largely driven by the large
temperature variability, which itself is dynamically driven, as discussed by
several authors (Siskind et al., 2015a; McCormack et al., 2006; Smith, 1995;
Froidevaux et al., 1989). Of interest here is that after 1 May the
interannual variability becomes very small, but is not zero. It also shows
that the relative abundance from year to year remains generally fixed
throughout the summer into the autumn. This small remaining difference is due
to chlorine chemistry, as seen below.
Altitude profiles of monthly and daily averaged ozone loss rates
from WACCM/NOGAPS for June 2009 (solid) and June 2008 (dashed) at
75∘ N.
Figure 6 shows the zonal and monthly averaged odd oxygen loss rates from the
HOx, ClOx and NOx catalytic cycles for June 2008 and 2009 at
75∘ N calculated by WACCM/NOGAPS. The expressions for these terms
are from Eq. (A1) of McCormack et al. (2006). The figure shows that the
contribution to total odd oxygen loss from chlorine chemistry maximizes in a
narrow layer from 1 to 3 hPa and that it is greater in 2009 than in 2008. This
is consistent with the greater ClO observed by MLS in 2009 as shown in
Fig. 3. The HOx cycle shows little change, but the NOx cycle actually
shows the opposite effect, i.e., decreased loss in 2009. This is perhaps
surprising and is worth documenting. Figure 7 shows the monthly averaged
NOx (i.e., NO + NO2) from WACCM for June for 75∘ N for 2009
and 2008. Above the stratosphere, from 1 to 0.1 hPa, NOx was higher in
2009. This is likely a legacy of enhanced descent from the upper mesosphere
observed earlier that spring. However, as discussed by Siskind et al. (2015b)
and also by Salmi et al. (2011) in their study of data from the Atmospheric
Chemistry Experiment Fourier Transform Spectrometer, there is no evidence
that these enhancements penetrated down to altitudes where the NOx
catalytic cycle affects ozone. Although SOFIE does not measure NO2, the
excellent agreement between WACCM NO and SOFIE NO documented by Siskind et
al. (2015b) gives us confidence that the WACCM NOx results are an accurate
reflection of reality. We suggest that the lower NOx from 1 to 8 hPa in
2009 is a legacy of greater winter/spring descent from the region of the NO
minimum in the mesosphere near 60–75 km.
Monthly averaged WACCM/NOGAPS NOx (= NO + NO2 for June
2009 (solid) and 2008 (dashed) at 75∘ N.
Percent change in monthly averaged ozone for June 2009 minus
June 2008 at 75∘ N. The solid line is from WACCM/NOGAPS and the
dashed line with stars is from MLS data.
Thus while there is some offsetting of the changes in the chlorine cycle by
the lower 2009 NOx, the net effect is that in the 1–2 hPa layer, the
overall odd oxygen loss is greater in 2009. Between 3 and 7 hPa, it is less in
2009. These changes agree well with observed ozone changes, as seen by MLS.
This is shown in Fig. 8 which presents an altitude profile of the ozone
change from WACCM/NOGAPS compared with MLS for June at 75∘ N. The
figure shows the relative 2009 ozone decrease near 1–2 hPa, corresponding
to the increase in chlorine loss. From 4 to 6 hPa, there is a small ozone
increase in 2009 which corresponds to the small reduction in NOx loss
suggested by Figs. 6 and 7.
Scatter plot of August monthly mean MLS O3 vs.
(a) MLS ClO and (b) SOFIE CH4 at 1.47 hPa. The
latitudes are near 78∘ N, corresponding to the latitude of the SOFIE
occultations in August.
Figure 9 shows that the ozone change over the entire 7-year period is
consistent with the above analysis for 2008 and 2009. Figure 9 presents
monthly averaged correlation coefficients between MLS ozone and MLS ClO
(Fig. 8a) and between MLS ozone and SOFIE CH4 (Fig. 8b) for 1.47 hPa.
Figure 9a shows that the approximate 5 % spread in ozone values is almost
perfectly anticorrelated with the 50 % ClO changes shown in Fig. 3.
Further, since we have previously shown that the summer ClO in the upper
stratosphere reflects the interannual variability in CH4, it is no
surprise that MLS O3, sampled at SOFIE latitudes, should almost perfectly
correlate with SOFIE CH4. This is shown in Fig. 9b.
Altitude profiles of linear correlation coefficients for SOFIE CH4 and MLS O3 (sampled at the SOFIE occultation latitudes).
The four curves are taken from zonal mean averages for May (solid), June (long dashes), July (dotted) and August (dotted/dashed).
Finally, Fig. 10 plots the linear correlation coefficient of CH4 and O3
as a function of altitude. Four curves are shown, corresponding to the four
monthly averages presented in Fig. 5. The figure shows that the correlation
maximizes in the 1–2 hPa region with values near and above 0.9. This is to
be expected from the chlorine cycle as shown in Fig. 6 above. Below
2–3 hPa, the NOx cycle becomes more dominant and the link to CH4
disappears. Thus the effects of uninterrupted wintertime descent of
mesospheric air on ozone may fall into two categories, separated by altitude.
From 1 to 2 hPa the ozone reductions result from chlorine enhancements; for
higher pressures, the potential for NOx enhancements dominates, provided
such enhancements were to make it down to those pressures.
Conclusions
We have shown how the chemical composition in the summertime upper stratosphere
depends upon dynamical activity from the previous winter and spring. Our main
result is to identify a new mechanism for summertime ClO and O3
variability, namely due to CH4 variations which, in turn, depend upon both
the magnitude of wintertime mesospheric descent and springtime planetary
waves. In 2009, prolonged mesospheric descent and a relative absence of
springtime wave activity lead to relatively low values of CH4 which
persisted throughout the summer. At the other extreme, in 2011, the lack of
strong winter descent combined with an intense frozen-in anticyclone event in
early April led to CH4 values which were more than twice that in 2009.
The excellent anticorrelation between MLS ClO and SOFIE CH4 both validates
our understanding of reactive chlorine partitioning and also offers a
framework for interpreting future observations. Due to orbital precession,
the latitudes of the SOFIE occultations have drifted away from polar region
and SOFIE is presently unable to monitor wintertime tracer descent. However,
based upon the results in this paper, perhaps MLS ClO data can be used as a
proxy for this. It would also be interesting to consider whether these
variations in ClO have any impact on O3 trend assessments. Both the strong
winter descent and the spring FrIAC phenomenon seem to be more common in
recent years . In principle, the enhanced
variability we have shown here might have to be considered, at least for trend
studies at high latitudes. Recent estimates of ClO trends
have only considered the tropics.
Our work shows that these CH4 and ClO variations have caused up to a 5 %
variation in upper stratospheric ozone throughout the summer and early fall.
This confirms the general role of chlorine chemistry in upper stratospheric
ozone. This also represents a second mechanism, in addition to that
associated with a descent of enhanced mesospheric NOx, by which descent of
mesospheric air can cause ozone reductions. Studies of spring- and summertime
ozone loss following strong descent years should take care to distinguish
between these two mechanisms. One way to distinguish them may be according to
altitude. Thus ozone decreases for p<3 hPa (z>40 km) are more
likely the result of low CH4, whereas for p>3 hPa (z<40 km),
NOx enhancements would dominate. A likely example of this second case is
shown in Fig. 1 of Randall et al. (2005).
Finally, the question of whether this variability would influence trend
analyses may be worth considering. There was earlier work using Upper
Atmospheric Research Satellite data to look at hemispheric differences in
ozone trends ; in light of the more recent dynamical
variability seen in the NH, and its now documented impact on ozone, perhaps
this should be revisited.