Introduction
The slowdown in the O3 decline and the beginnings of recovery of the
ozone layer have been documented (Newchurch et al., 2003; Yang et al., 2006). Monitoring of the ozone
layer continues to be critical in order to understand ozone recovery as the
CFC (chloroflourocarbon) burden in the stratosphere decreases. A number of observational studies
have quantified the global distribution of changes to the O3 layer and
revealed distinct patterns and variability which show that O3 trends
are not spatially uniform. One consistent result is that over decadal time
scales, equatorial O3 in a vertical layer near 30 km (corresponding to
∼ 10 hPa) often varies very differently from O3 in the
rest of the middle to upper stratosphere. Kryölä et al. (2013), using measurements from
the Stratospheric Aerosol and Gas Experiment (SAGE) from 1984 to 1997, show a
general decrease in O3 which is statistically significant over much of
the stratosphere, but an increase in equatorial O3 (albeit not
statistically significant) in the 30–35 km region. Conversely, for the
period 1997–2011 Kryölä et al. (2013) show a general increase in O3 from SAGE and
Global Ozone Monitoring by Occultation of Stars (GOMOS) measurements, but a
statistically significant decrease near 30 km in the tropics. Bourassa et al. (2014)
combine SAGE measurements with measurements from the Optical Spectrograph
and InfraRed Imager System (OSIRIS) instrument and, again splitting the data
into pre- and post-1997 periods, find very similar results. Damadeo et al. (2014) compute
a SAGE trend from 1998 to 2005 and find a positive trend near 30 km in the
tropics and Northern Hemisphere, but a negative trend in the Southern
Hemisphere at this level. Measurements from the Scanning Imaging Absorption
Spectrometer for Atmospheric Chartography (SCHIAMACHY) instrument for the
period 2002–2012, reported by Gebhardt et al. (2014), show a pattern similar to the
1997–2011 pattern reported by Kryölä et al. (2013), i.e., a strong statistically
significant decrease in tropical O3 in the 30–35 km region, while most
of the middle atmosphere shows a slight increase in O3. Finally,
Eckert et al. (2014) using Michelson Interferometer for Passive Atmospheric Sounding
(MIPAS) data from 2002 to 2012, also show a general increase in O3 in most
regions, but find statistically significant negative trends in the tropics
from ∼ 25 to 5 hPa. Eckert et al. (2014) note that increased upwelling
has been suggested as an explanation for ozone decreases, but, in referring
to these trends, they conclude that “upwelling does not provide a
sufficient explanation for the negative values in the tropical
mid-stratosphere.”
Ozone at 10 hPa over the equator is particularly sensitive to catalytic
cycles involving the odd nitrogen (NOy) chemical family (Olsen et al., 2001;
Brasseur and Solomon, 1986). Ravishankara et al. (2009) showed that N2O would be the dominant ozone depleting
substance emitted in the 21st century, and pointed out that nitrogen
oxides contribute most to O3 depletion just above where the O3
mixing ratios are the largest. Portmann et al. (2012) calculated the effects of a surface
boundary increase of 20 ppbv of N2O (an increase expected over
∼ 20 years in the IPCC A1B scenario) on O3. They showed
that this increase in N2O emission would lead to a global mean decrease
of ∼ 0.5–0.7 % in O3 just above the peak of the ozone
mixing ratio (0.1 DU km-1 in the 30–35 km region where O3 has a
density of ∼ 15–20 DU km-1 based on their Fig. 2). In
mixing ratio terms, this gives a rate of ∼ 5–7 ppbv yr-1.
Plummer et al. (2010) studied O3 changes in a model including GHGs (greenhouse gases) and ODSs (ozone-depleting substances). They ran
two experiments with a faster Brewer–Dobson circulation, and these two
experiments showed, at 10 hPa in the tropics, a decrease in reactive
nitrogen and an increase in both O3 and N2O at 10 hPa relative to
the experiments with a slower circulation. Thus variations in O3 and
N2O at 10 hPa can be either correlated or anti-correlated depending
upon whether they are driven primarily by circulation or by changes in
N2O entering the stratosphere.
In addition to long-term anthropogenically driven changes, events such as
the eruption of Mt. Pinatubo may alter the chemistry and dynamics of the
stratosphere for extended periods. Aquila et al. (2013) compared a reference model with a
model which simulated the effect of the volcanic aerosols on both chemistry
and dynamics. They calculated an increase in O3 of ∼ 2 % at 10 hPa in the tropics slightly more than a year after the eruption,
with no strong latitudinal variation. Damadeo et al. (2014) attempt to disentangle
anthropogenically driven changes from Pinatubo eruption driven changes using
an aerosol based volcanic proxy.
Previous observational work has correlated O3 interannual variability
in the tropics near 10 hPa with changes in specific odd nitrogen compounds;
however, these studies were only for relatively short time periods compared
with the O3 studies referenced above. Randel et al. (2000) showed that the Halogen
Occultation Experiment (HALOE) observed increasing NO + NO2 coincident
with decreasing O3 from 1992 to 1997, but that these variations
leveled-off during the last years of HALOE measurements which were then
available (1998–2000). The rate of O3 decrease from 1992 to 1996 was
faster than 100 ppbv yr-1 just above 10 hPa in the tropics. The HALOE
measurements of NO2 at ∼ 10 hPa from 1993 to 1997 were shown
to be consistent with a decrease in upward transport (Nedoluha et al., 1998) and increased
photolysis of N2O, the source of stratospheric NOy.
The present study extends the previous observational studies with a
combination of 21 years of ozone data from the UARS (Upper Atmosphere Research Satellite) HALOE and the Aura
Microwave Limb Sounder (MLS) measurements, plus nitrogen species data from
HALOE, MLS and the Atmospheric Chemistry Experiment (ACE). Our results
confirm the existence of the 10 hPa tropical ozone trend anomaly and link it
to a correspondingly consistent local change in the nitrogen species which
affect O3. The resulting rate of change in O3 and in the nitrogen
species is an order-of-magnitude faster than changes predicted from model
calculations based upon changes in anthropogenic emissions.
Measurements from HALOE, aura MLS, and ACE
We make use of measurements from the HALOE, MLS, and the Fourier transform
spectrometer measurements from ACE. HALOE measurements of O3, NO, and
NO2 are available from 1991 to 2005. HALOE used the solar occultation
technique which provided ∼ 28–30 profiles per day in two
latitude bands, one at sunrise and one at sunset. The latitude bands drifted
daily so that near global latitudinal coverage was provided in both sunrise
and sunset modes five times over the course of a year. The trends in the
HALOE O3 measurements have been compared against SAGE II (Nazaryan et al., 2005) and
differences have been found to be on the order of less than 0.3 % per year
in a majority of latitude bands at 25, 35, 45, and 55 km.
MLS measurements of O3 and N2O are available since 2004. MLS
measurements are available over a global range of latitudes on a daily
basis. The stratospheric O3 product has been validated by Froidevaux et al., (2008). The
N2O measurements have been validated by Lambert et al., (2007).
Since 2004 ACE has been measuring O3, N2O, and the nitrogen
species that constitute the bulk of NOy (NO, NO2, HNO3, and
N2O5). As a solar occultation instrument it, like HALOE, provides
∼ 28–30 profiles per day in two latitude bands, one at sunrise
and one at sunset. The ACE O3 measurements have been validated by
Dupuy et al. (2009), and the NO and NO2 measurements were validated by
Kerzenmacher et al. (2008).
The solar cycle and linear trend calculations
In cases where species are affected by the solar cycle, one of the
challenges in interpreting decadal scale trends in the stratosphere is
separating these trends from solar cycle induced variations. Model studies
provide some guidance as to the expected solar cycle variations in the
species of interest. Egorova et al. (2005) used the SOCOL Chemistry Climate Model (CCM)
and found that at 30 km O3 was higher at solar maximum when compared to
solar minimum, but that the difference was < 3 %. The N2O
mixing ratio at 30 km from 30∘ S to 30∘ N was found to be no more than
2 % higher at solar minimum compared to solar maximum, and no more 4 %
from 30 to 60∘ N and 30 to 60∘ S. Schmidt et al. (2010) used
the HAMMONIA general circulation and chemistry model, and found an
equatorial O3 sensitivity of ∼ 1.4 ± 0.4 % 100-1 solar
flux units (sfu), where the difference between the 1989 solar max and the
1986 solar min is 166 sfu. The study of Remsberg and Lingenfelser (2010) shows a 3 % ozone
maximum–minimum response to the solar cycle at ∼ 35 km
(∼ 7 hPa) from the SAGE II measurements, with results from the
HALOE measurements and from model calculations showing a smaller ozone
response to the solar cycle. As discussed in Hood and Soukharev (2006), NOy in the upper
stratosphere is also affected by the solar cycle. They place an upper limit
of ∼ 10 % on the solar cycle variations in NOy in the
tropical mid-stratosphere. The model calculations in Egorova et al. (2005) show an
NO2 solar cycle variation of < 1 %, and Nedoluha et al. (1998) show a
similarly small variation from the CHEM2D model (Bacmeister et al., 1998).
Throughout this study we will calculate trends based on a function including
terms to fit the annual, semi-annual, QBO (quasi-biennial oscillation), plus a constant term and a linear
trend term. The QBO terms were calculated using the Center for Climate
Prediction 30 and 50 hPa winds anomalies obtained from www.cpc.ncep.noaa.gov/data/indices/. In addition to these terms, we have
calculated trends from the HALOE measurements both with and without the
inclusion of a solar cycle term, where the solar cycle fit is calculated
using the Mg II values obtained from the Laboratory for Atmospheric and
Space Physics (LASP) Interactive Solar Irradiance Data Center at
lasp.colorado.edu/lisird. We will only show HALOE trend calculations where a
solar cycle term has been included, but we have compared trends with and
without the solar cycle term and found that the results are similar.
The MLS measurement time series used here extends from 2004 to 2014, and
therefore clearly does not extend over a full solar cycle. The linear trend
calculations from MLS measurements which will be shown cover the period
August 2004 to May 2013. Because solar cycle 24 is particularly weak, the Mg
II values in 2013 are comparable to those in 2004, so solar effects are
unlikely to cause a trend in the MLS data set used here. In order to provide
an estimate of the uncertainty in the trend, which is introduced by the
presence of a solar cycle, we will show some MLS results both with and
without the inclusion of a solar cycle term. We will show that in the region
of greatest interest, near the tropics at ∼ 10 hPa, the MLS
trends appear to be nearly insensitive to the presence of a solar cycle.
Measurements of O<inline-formula><mml:math display="inline"><mml:msub><mml:mi/><mml:mn mathvariant="normal">3</mml:mn></mml:msub></mml:math></inline-formula>, 1991–2014
In Fig. 1 we show the annual median O3 anomalies from
5∘ S to 5∘ N as measured by both HALOE and Aura MLS at 10 hPa. This
figure also shows that the O3 decrease at 10 hPa has occurred gradually
over the period shown. There have been numerous studies combining O3
time series from multiple satellites to derive long-term trends (e.g., Jones et al., 2009;
Kryölä et al., 2013), and there are several projects underway to provide long-term data
records of stratospheric composition, so we will not attempt here to produce
a combined HALOE-MLS O3 time series for trend calculations. The MLS
time series anomalies shown have simply been offset by a shift in mixing
ratio so that the anomaly point for 2005 (which covers data taken during the
period July 2004 through June 2005) agrees with the HALOE anomaly at that
point. The anomalies are calculated by fitting annual and semi-annual cycles
to each data set separately and then calculating annual median differences
from this fit. The annual anomaly is sampled four times per year so that
each point represents an anomaly over either January–December, April–March,
July–June, or October–September. Each measurement is therefore included in
four data points in the figure. Having removed the annual cycle, the primary
variation in O3 in this region is caused by the QBO. In addition to
these QBO variations, there is a clear decrease in O3 over the 21 years
shown.
Annual median ozone anomalies at 10 hPa
5∘ S–5∘ N from HALOE (green; HALOE data is
actually shown on its native grid at 30 km, which is ∼ 10 hPa)
and MLS (red). Annual anomalies are shown four times per year; hence each
measurement is included in four data points. The MLS anomalies have been
shifted by a constant mixing ratio so that the HALOE and MLS annual
anomalies for 2005 (covering the period July 2004–June 2005) agree. The MLS
data from July 2013 onwards is indicated dashed in order to indicate that
this data will not be used in any of the linear trend calculations to be
shown.
![]()
An estimate of the uncertainty in these annual medians can be obtained from
the standard deviation of the individual anomalies. The average value of
σ n-1/2 for the annual median HALOE O3 anomalies is
0.026 ppmv. The last annual anomaly has the largest uncertainty with σ n-1/2= 0.056 ppmv. For the MLS, which has many more measurements, the
largest annual median uncertainty calculated by this method is 0.0027 ppmv.
In Fig. 2 we show the linear trend in the global HALOE ozone measurements
from 1991 to 2005. Remsberg (2008) show a quite similar figure (their Fig. 13) for linear trends in HALOE O3, but in % decade-1. They found that
the trends near 10 hPa from ∼ 25∘ S to ∼ 25∘ N had a confidence interval of > 90 % for their partial
tank order correlations (Remsberg et al., 2001). The trend is negative (i.e.,
O3 is decreasing) near ∼ 10 hPa with the most negative
values occurring in the tropics. Most of this study will focus primarily on
the causes of this O3 decrease in this region. In general the results
are very similar whether or not a solar cycle is included in the fit, but
the local tropical minimum at ∼ 4 hPa does not appear when
such a term is not included.
The calculated linear trend in HALOE ozone for 1991–2005. The HALOE data
has been sorted into 11 individual 10∘ latitude bins from
55∘ S to 55∘ N. Regions where the magnitude of the
trend is < 0.01 ppmv year-1 are indicated in white.
![]()
HALOE measurements ceased in 2005, and Aura MLS has been providing O3
measurements since 2004. In Fig. 3 we show the linear trend in O3 as
measured by MLS. MLS shows that the negative ozone trend in the tropics near
∼ 10 hPa continued from August 2004 to June 2013. Inclusion of
the most recent MLS data (from July 2013 to September 2014) does not result in a
qualitative change in Fig. 3, but does reduce the magnitude of the
measured trends. Again, the O3 linear trends shown in Fig. 3 have
been calculated with a solar cycle included in the fit, but the results are
very similar with and without a solar cycle term. Several other data sets
have also shown decreasing O3 near 10 hPa in the tropics. Kryölä et al. (2013) has
shown a decrease for 1997–2011 from SAGE and GOMOS, Gebhardt et al. (2014) for 2002–2012
using measurements from SCIAMACHY, and Eckert et al. (2014) for 2002–2012 using MIPAS
measurements. There is some overlap between the negative HALOE O3 trend
(1991–2005) and the positive SAGE O3 trend shown by (Kryölä et al., 2013; 1984–1997). Given the excellent agreement between SAGE II and HALOE trends
(e.g., Nazaryan et al., 2005), and the eruption of Mt. Pinatubo near the middle of the
1984–1997 time series, we expect this difference between the 1984–1997 and
1991–2005 trends is caused by a real change in O3 trends in the
tropical 10 hPa region in between 1991 and 1997.
The O3 linear trend calculated from MLS data from August
2004 to May 2013. Contour lines are shown at ±0.01, 0.02, 0.03, 0.04,
0.06, 0.08 ppmv year-1. The MLS data has been sorted into twelve
10∘ latitude bins from 60∘ S to
60∘ N. Regions where the magnitude of the trend is < 0.01 ppmv year-1 are indicated in white.
![]()
While Fig. 1 shows a general decrease in O3 at 10 hPa over the entire
HALOE measurement period, and Fig. 3 shows that this trend continued into
the MLS measurements period, such trends do not always persist over such
extended periods. For example, away from the tropics, the 1991–2005 HALOE and
2004–2013 MLS trends near 10 hPa show clear, hemispherically dependent,
differences. Whereas the HALOE trends show a decrease in O3 at all
latitudes near 10 hPa, the MLS trends show a sharp increase in O3 at
southern mid-latitudes, and a smaller decrease at similar pressure levels in
northern mid-latitudes.
Just as the HALOE and MLS trends show clear differences away from the
tropics, they also show clear differences in the tropics at other levels.
The 1991–2005 HALOE trend shows an increase in tropical O3 near 30 hPa,
but this is dominated by the strong increase from ∼ 1991 to 1999,
followed by a period of stability in this region from 1999 to 2005. The MLS
O3 measurements suggest that this period of stability near 30 hPa
continues through 2013. However Gebhardt et al. (2014) do show statistically significant
O3 increases in the 2002–2012 SCIAMACHY measurements below 30 km with
two local maxima, one near 22 km and one near 27 km, while Eckert et al. (2014) show an
O3 increase from 2002 to 2012 near 22 km (∼ 50 hPa) but a
decrease near 27 km (∼ 25 hPa) from the MIPAS measurements. In
their 1984–1997 O3 trends Kryölä et al. (2013) show an increase at 24 km, but a much
larger decrease at 21 km. Thus, while several measurements show decadal
scale tropical trends near 10 hPa, such trends to not appear to be common
near 30 hPa nor at other latitudes near 10 hPa.
The effect of changes in nitrogen species on ozone
As noted above, O3 in the tropical mid-stratosphere is particularly
sensitive to changes in NOy which result from photodissociation and
oxidation of N2O (Olsen et al., 2001), and long-term increases in anthropogenic N2O
emission are expected to play a significant role in causing future
decreases in O3 (Portman et al., 2012). However, N2O is also a sensitive indicator of
upward transport and, as we show below, these variations in transport lead
to a positive, not negative correlation between N2O and O3. Figure 4 presents the correlation between MLS N2O and O3 from 2004 to 2013.
These correlations are calculated by first finding a zonal monthly median
for each year of MLS data and then subtracting from each of these the
average MLS monthly median for that month. Note the strongly positive
correlation precisely where the observed long-term trends indicate ozone
decreases. Figure 5 presents monthly median MLS N2O and O3 data
from 5∘ S–5∘ N at 10 hPa. The positive correlation between N2O
and O3 is clearly present on seasonal and interannual timescales and
rules out an anthropogenic increase in N2O as the cause of the long-term ozone decreases we observe.
Correlation coefficients between N2O and O3
calculated from monthly median anomalies from MLS data as a function of
latitude and pressure. Results are shown for regions where the correlation
(or anti-correlation) is > 0.6.
![]()
This positive correlation between N2O and O3 in the tropical
middle stratosphere can be readily understood in the context of the
relationship between N2O, NOy and O3. This is demonstrated in
Fig. 6 which presents ACE measurements of O3, N2O and the
species which make up the bulk of NOy at 30 km (∼ 10 hPa)
from 10∘ S–10∘ N. While ACE does not provide the daily measurement
coverage in the tropics obtained by MLS, it does measure all of the species
relevant to the nitrogen chemistry which determines O3 near
∼ 10 hPa in the tropics. Like MLS, ACE shows a strong positive
correlation between N2O and ozone. ACE also shows the expected
anticorrelation resulting from the chemistry of O3 and NOy. Figure 6c shows the anti-correlation between NOy and N2O without which
the correlation between N2O and O3 would not exist. This
anti-correlation of N2O and NOy can be understood as a coupled
chemical/dynamical effect. During periods when upward transport is slower,
more N2O at a given altitude is dissociated, thus producing more
NOy at that altitude. We thus conclude that over the period of the
MLS measurements, the effect of changes in transport on N2O in this
region on NOy and hence O3 dominate any increase in N2O due
to changing tropospheric emissions.
Monthly median N2O (red) and O3 (blue) mixing
ratios at 10 hPa from MLS measurements between 5∘ S and
5∘ N.
![]()
As indicated in Sect. 2.1, the MLS instrument has been operational for
less than a full solar cycle; hence for tropical trend calculations we show
results both with and without the inclusion of a solar cycle term. In Fig. 7 we show the calculated profiles as a function of pressure as derived from
eight (constant term, two annual terms, two semi-annual terms, two QBO
terms, and a linear trend) and nine (including a solar cycle) parameter fits
to the monthly median MLS measurements. Figure 7 shows the profiles (the
constant terms from the fits) in addition to the linear trend and the net
effect of 8 years of such a trend (2004–2005 vs. 2012–2013). The O3 trend
results are in good agreement with those shown by Gebhardt et al. (2014) for August
2004–April 2012, where the fastest decreasing trend in MLS O3 is
∼ 7 % decade-1. Gebhardt et al. (2014) show that the MLS trends in O3
are not statistically different from those observed by SCIAMACHY or OSIRIS.
While the inclusion of the solar cycle term fit clearly does affect the
linear trend at some levels, it does not alter the qualitative result that
O3 and N2O both show a statistically significant decrease over a
similar range of pressures near 10 hPa. This further reinforces the
conclusion that this decrease in O3 is caused by an increase in
NOy (resulting from increased dissociation of N2O) during this
period.
ACE measurements of O3, N2O, and the key
members of the NOy family,
NO + NO2+ HNO3+ 2 × N2O5.
Measurements are shown for 10∘ S–10∘ N at 30 km.
Both sunrise and sunset measurements are included.
![]()
Left hand panel: annual average MLS profiles of O3 (blue;
top scale) and N2O (red; bottom scale) from
5∘ S–5∘ N. Shown are the constant term derived
from the fit to the August 2004–May 2013 MLS measurements (solid), and the
same term with an added 8-year shift (thus approximating the difference
between the 2004–2005 and 2012–2013 MLS annual average) based on the linear trend
applied over a period comparable to the length of the MLS data set (dotted).
Right hand panel: linear annual trend calculated with a solar cycle term
included (solid) and without a solar cycle term (dashed). Error bars
(2σ) are similar for fits with and without the solar cycle, and are
shown only for the former.
![]()
As was shown in Fig. 2, HALOE measurements showed a decrease in O3
from 1992–2005 at 10 hPa from 5∘ S–5∘ N. While HALOE did not provide
measurements of N2O, and did not provide the full complement of
NOy species that is available from ACE, it did provide measurements of
two of the key odd-nitrogen species, NO and NO2.
In Fig. 8 we show annual median HALOE anomalies in O3 alongside those
of NO + NO2. Because NO + NO2 has a strong diurnal component
(unlike the set of NOy measurements provided by ACE), the anomalies for
both species are calculated separately for sunrise and sunset measurements.
We have multiplied the sunset NO + NO2 measurements by 0.4 so that they
fit onto the same scale as the sunrise measurements. The average (maximum)
σ n-1/2 value for sunrise NO + NO2 is 0.064 ppbv
(0.082 ppbv), while for the sunset measurements (before multiplication by 0.4) it
is 0.122 ppbv (0.22 ppbv). For the O3 sunrise measurements the average
(maximum) σ n-1/2 value is 0.045 ppmv (0.064 ppmv), while for
the sunset measurements it is 0.049 ppmv (0.081 ppmv).
Figure 8 shows that NO + NO2 generally was increasing over the course
of the HALOE measurements and that this increase tracked the ozone decrease,
both on a year-to-year timescale (dominated by the quasi-biennial
oscillation, QBO) and over the full 1992–2005 time period. There is a slight
(∼ 3-month) apparent phase-lag between the sunset and sunrise
NO + NO2 measurements from ∼ 1998 to 2003, which is not
apparent in the O3 anomalies and for which we have no explanation. With
the exception of this feature, the general consistency between the QBO
driven variations in O3 and NO + NO2, and the trend which is
apparent in both the O3 and NO + NO2 measurements, provides added
confidence that the decrease in O3 and the increase in NO + NO2
measured by HALOE from 1992 to 2005 are both correct and, further, are coupled.
As we concluded from the MLS measurements from 2004 to 2013, this change
suggests a slowdown in upward transport in this region from 1992 to 2005. Note,
this is consistent with the results of Nedoluha et al. (1998) who interpreted the
decreases in upper stratospheric CH4 from 1992 to 1996 as linked with a
simultaneous increase in NO2 at 30 km. Our results here extend that
early result to encompass the entire 13 year UARS mission.
Annual median HALOE O3 (top) and NO + NO2
(bottom) anomalies at 10 hPa from 5∘ S to 5∘ N.
Annual anomalies are shown four times per year; hence, each measurement is
included in four data points. Results are shown separately for sunrise
(solid) and sunset (dashed). Sunset NO + NO2 anomalies have
been multiplied by 0.4 so that they fit on the same scale as the sunrise
anomalies.
![]()
In Fig. 9 we show the calculated linear trends in the HALOE O3 and
NO + NO2 measurements. As in Fig. 8 we separate the HALOE sunrise and
sunset measurements, and calculate trends for four separate measurements:
sunrise and sunset O3 and sunrise and sunset NO + NO2.
Encouragingly, despite having very different vertical profiles, the shape of
the sunrise and sunset NO + NO2 trend profiles are very similar. The
O3 sunrise and sunset trends also agree well, and the pressure level of
the minimum of the observed decrease in these O3 measurements
corresponds closely with the maximum in the observed increase in the
NO + NO2 measurements.
Model calculations
In order to better understand the changes in the observed species we have
employed the two-dimensional chemical transport model (CHEM2D; Bacmeister et al., 1998). The
model includes parameterized gravity wave and planetary wave drag and is
ideal for understanding tracer transport and the response of the global
middle atmospheric circulation to external forcings. Compared with those
earlier studies, the present model has an improved vertical resolution (1
instead of 2 km). CHEM2D's most recent applications have included simulating
the solar cycle variations of polar mesospheric clouds (Siskind et al., 2005) and studying
the response of stratospheric ozone to both the solar cycle and the tropical
quasi-biennial oscillation (McCormack et al., 2007).
Left hand panel: annual average HALOE profiles of O3 at
local sunset (blue; top scale), local sunrise (black; top scale) and
NO + NO2 at local sunset (red; bottom scale), and local
sunrise (orange; bottom scale) from 5∘ S to 5∘ N.
Results are shown for the first year of HALOE measurements (solid) and with
an 8-year shift (to allow for comparison with Fig. 7) using the annual
average trend shown on the right-hand panel (dashed). Right hand panel:
1-year changes based on linear trends over
5∘ S–5∘ N calculated from 1991 to 2005. Line colors
are the same as in the left-hand panel.
![]()
We will show results from two model runs, each of which has been integrated
for 12 years to ensure stability from year-to-year. Since the goal of the
model was to test whether dynamical changes would affect N2O, NOy,
and O3 at the equator, we introduced a very simple perturbation. The
two models differ only in that, in one case, we added a small heat source of
0.3 K day-1, centered at 18 km at the equator, similar to Experiment 7 of
Bacmeister et al. (1998). In addition, we recognize that the model differences shown
represent two equilibrium solutions while the calculated trends show the
effects of an atmosphere changing over time. Nonetheless, a comparison of
these two models can serve as an indication whether it is possible to
reproduce the observed changes in the measured species with a dynamical
perturbation.
Figure 10 shows the change in N2O, NOy, and O3 at the equator
resulting from this dynamical perturbation. The absolute values for these
three species at 10 hPa are in very good general agreement with those shown
in Fig. 6 from the ACE measurements. The N2O chemistry is relatively
simple, and N2O at all levels is lower for the case of the slower
tropical ascent which offers more time for dissociation. At 10 hPa the case
with the slower ascent shows ∼ 22 ppbv less N2O. Unlike
the measurements, however, the N2O in the model changes over a deep
layer, covering the entire pressure range from 50-1hPa.
Annual average altitude profiles of O3, NOy,
and N2O for the equator. The solid curve is a baseline run,
while the dashed curve is a simulation which includes an additional 0.3 K day-1 heat source in the lowermost stratosphere which acts to increase the
tropical upwelling.
![]()
The calculated equatorial N2O changes shown in Fig. 10 correspond
with calculated O3 and NOy profile changes which are in the same
sense and general magnitude as the observations. Thus the calculation with
lower N2O yields increased NOy due to increased oxidation. via
N2O + O(1D) - > 2NO; the baseline model with ∼ 22 ppbv less N2O shows an increase of ∼ 1.3 ppbv in
NOy, so ΔNOy/ΔN2O ∼ 0.065. This
is similar to the ΔNOy/ΔN2O in the ACE
measurements in Fig. 6c, which is ∼ 0.75. Regarding ozone,
the baseline model with less heating and 1.3 ppbv greater NOy shows
about ∼ 0.26 ppmv less O3 at 10 hPa. This yields a
ΔO3/ΔNOy of ∼ 200 which is on the
order of, but somewhat less than, the observed ΔO3/ΔNOy from the ACE measurements (which is closer to ∼ 330). We can also directly compare the calculated quantity ΔO3/ΔN2O in the model and in the ACE and MLS measurements;
here we also get a somewhat reduced ozone response compared with the
observations. Whereas the model shows a ΔO3/ΔN2O
of ∼ 12, ACE and MLS both show ΔO3/ΔN2O ∼ 25). About 25 % of this difference is because
the model ΔNOy/ΔN2O sensitivity is slightly less
than observed. Also, since our approach towards diurnal averaging requires a
specific of a single night-to-day ratio (cf. Summers et al., 1997), it is possible that we
are underestimating the diurnally averaged response of O3 to NOx
chemistry, which varies strongly with the time of day. Nevertheless, the
qualitative agreement between the model and the ACE and MLS measurements
supports the idea that the observed O3 change can be caused by a
dynamical perturbation.
While the model runs do support the suggestion that the changes in O3
and N2O observed near the equator at ∼ 10 hPa can be
caused by a dynamical perturbation, we do note that this particular
dynamical perturbation also shows large differences in other regions where
the measured trends are small and/or vary in a temporally different manner
than do the tropical 10 hPa measurements. No doubt a number of dynamical
changes affected N2O over the period 1991–2013, and these variations
drove changes in NOx and in turn O3. What we conclude here is
that, because of changes in transport, the N2O which arrived in this
region experienced significantly more dissociation in 2013 than in 2004,
and, based on inferences from the HALOE O3 and NOx measurements,
that this trend was also present throughout much of the HALOE measurement
period. We note that there is a burgeoning literature debating the
possibility of changes in the stratospheric circulation (cf. Butchart, 2014 and
references therein). The results presented here may serve as a useful
constraint for these analyses of long-term stratospheric variability.