Introduction
In the stratosphere, the photo-oxidation of N2O is the main source of
all nitrogen species (NOy). About 97 % of the stratospheric
NOy budget can be explained by NO, NO2, HNO3,
ClONO2, and N2O5 compounds, and the partitioning between
reactive and reservoir nitrogen species is an important issue in
stratospheric ozone chemistry (e.g. Wetzel et al., 2002; Brohede et al.,
2008). Nitrogen oxides (NOx= NO + NO2) are major catalysts
responsible for significant ozone destruction in the middle stratosphere. In
the gas phase, NOx interacts with the hydrogen and halogen species in
catalytic cycles affecting ozone loss rates in the lower stratosphere (e.g.
Portmann et al., 1999; Salawitch et al., 2005). Therefore, NOx can also
buffer the ozone destruction by halogenated compounds through the formation
of ClONO2 and BrONO2 (e.g. Rivière et al., 2004). The
HNO3 reservoir is formed from NOx indirectly via the hydrolysis of
N2O5 on liquid sulfate aerosols:
N2O5+H2O(aq)→2HNO3.
It has been shown that models need to include Reaction (R1) to better
reproduce observations of NOy partitioning at midlatitude for
background aerosol conditions (i.e. in volcanically quiescent periods) in
the lower stratosphere (Rodriguez et al., 1991; Granier and Brasseur, 1992;
Fahey et al., 1993; Webster et al., 1994; Salawitch et al., 1994b; Sen et
al., 1998). This reaction tends to decrease NOx amounts and reduces the
ozone loss efficiency associated with the NOx catalytic cycle as the
less reactive nitrogen reservoir HNO3 is formed (e.g. Rodriguez et al.,
1991; Weisenstein, 1991; McElroy et al., 1992). Reaction (R1) is fairly
insensitive to temperature and has the potential to greatly reduce reactive
nitrogen globally, even under background aerosol conditions.
The hydrolysis of ClONO2 can be expressed by
ClONO2+H2O(aq)→HNO3+HOCl.
It results in the additional formation of HNO3 on sulfate aerosols and
in the formation of reactive chlorine in sunlight, where HOCl is rapidly
photolysed releasing Cl radicals (e.g. Hofmann and Solomon, 1989; Prather,
1992; McElroy et al., 1992). This heterogeneous reaction is highly dependent
on the water content in the aerosols and has been shown to be of
considerable importance in determining the abundance of active chlorine
available to destroy ozone under some conditions, i.e. for temperatures
typically below 210–215 K and where HNO3 photolysis rates are slow
(typically in winter at high latitudes) (e.g. Hanson et al., 1994; Tie et
al., 1994; Borrmann et al., 1997). However, for higher temperatures the
ClONO2 hydrolysis is not expected to be significant enough to compete
with Reaction (R1) on the NOy partitioning under these conditions (Fahey
et al., 1993; Cox et al., 1994; Sen et al., 1998). Also, the reaction
ClONO2+HCl(aq)→HNO3+Cl2
of ClONO2 with dissolved HCl in sulfuric acid droplets has negligible
effects on chlorine activation at such temperatures (Hanson et al., 1994;
Borrmann et al., 1997).
Some works also suggest that the hydrolysis of BrONO2,
BrONO2+H2O(aq)→HNO3+HOBr,
on background sulfate aerosols also plays a significant role in ozone
depletion in the lower stratosphere with rates almost independent of
temperature, making this reaction efficient at all latitudes and for all
seasons (Hanson and Ravishankara, 1995; Hanson et al., 1996; Lary et al.,
1996; Randeniya et al., 1997; Erle et al., 1998).
After large volcanic eruptions, the aerosol loading in the stratosphere and
the surface area densities (hereafter SADs) available for Reaction (R1) to
occur are dramatically enhanced (e.g. Deshler et al., 2003). As a result,
the amount of ozone-depleting NOx is strongly reduced (e.g. Prather,
1992; Johnston et al., 1992; Fahey et al., 1993; Mills et al., 1993; Solomon et
al., 1994; Kondo et al., 1997; Sen et al., 1998; Dhomse et al., 2015),
whereas HNO3 amounts increase (Koike et al., 1993, 1994; Webster et al.,
1994; Rinsland et al., 2003) as shown for the Pinatubo
aerosols. Different chemical impacts on stratospheric ozone are expected
depending on the altitude. In the middle stratosphere (above
∼ 30 hPa), where ozone loss is dominated by NOx, the presence of volcanic
aerosols can result in layers of increased net production of ozone due to
the suppression of the NOx cycle by the N2O5 hydrolysis
(Hofmann et al., 1994; Bekki and Pyle, 1994; Tie and Brasseur, 1995). In the
lower stratosphere, halogen (ClOx and BrOx) and hydrogen
(HOx) radicals play a dominant role in ozone depletion, and their
abundances, which depend on NOx levels, are increased (in particular
for halogen species, as the rate of gas-phase conversion of ClO into the
ClONO2 reservoir is reduced), resulting in an enhanced catalysed ozone
loss (McElroy et al., 1992; Granier and Brasseur, 1992; Brasseur and
Granier, 1992; Hofmann et al., 1994; McGee et al., 1994; Bekki and Pyle,
1994; Salawitch et al., 1994a, 2005; Tie et al., 1994; Solomon et al., 1996;
Solomon, 1999).
However, the NOx-to-HNO3 conversion by Reaction (R1) shows
saturation as the aerosol SAD increases because the amount of N2O5
present in the stratosphere is limited by its production rate by the gaseous
reaction NO2+ NO3 (Fahey et al., 1993; Prather, 1992; Mills et al.,
1993; Tie et al., 1994; Solomon et al., 1996; Kondo et al., 1997; Sen et
al., 1998). Consequently, ozone loss rates are expected to be limited
because the saturation of the NOx / NOy response to the aerosol
increase dampens the increase in ClO / Cly (Fahey et al., 1993; Tie et
al., 1994). Reaction (R2) does not show such a rapid saturation resulting in
enhanced ozone depletion by chlorine catalytic cycles in cold air masses as
the aerosol loading increases (Fahey et al., 1993). The BrONO2
hydrolysis through Reaction (R4) is primarily dependent on the aerosol
loading and is enhanced in periods of high volcanic aerosol loading. The
resulting increase in BrOx and HOx radical concentrations and
decrease in HCl (due to enhanced OH) accompanied by an increase in ClOx
radicals is expected to give further ozone loss in the lower stratosphere at
all latitudes and seasons (Lary et al., 1996).
In periods following major eruptions, the year-to-year variability in
stratospheric ozone at northern midlatitudes appears closely linked to
dynamical changes induced by the volcanic aerosol radiative perturbation
(e.g. Telford et al., 2009; Aquila et al., 2013) and to changes in chlorine
partitioning (e.g. Solomon, 1999; Chipperfield, 1999). Effects on
stratospheric chemistry are expected in periods of elevated chlorine levels
from anthropogenic activities (Tie and Brasseur, 1995; Solomon et al.,
1996). In the past decade no event comparable to the 1991 Pinatubo or 1982
El Chichón eruptions was observed. However, several volcanic eruptions,
though of much lesser amplitude, impacted the aerosol burden in the lower
stratosphere over periods of months (Vernier et al., 2011). These
“moderate” eruptions have occurred in a period of still high chlorine loading with a potential impact on stratospheric ozone chemistry. Their effects depend on
the amount of released SO2 and on latitudes and altitudes of injection
which directly influence aerosol residence times. The season of the eruption
is also important for photochemical processes which are directly connected
to temperatures and solar illumination.
The goal of this paper is to show how such moderate eruptions are likely to
modify the chemical balance of the Northern Hemisphere lower stratosphere at
periods excluding wintertime or springtime halogen-activating photochemistry.
We specifically focus on the eruption of the Sarychev volcano on 15 and
16 June 2009, which injected 0.9 Tg of sulfur dioxide into the lower stratosphere
(Clarisse et al., 2012), resulting in enhanced sulfate aerosol loading up to
19 km, for a period of about 8 months ending before winter (Haywood et al.,
2010; Kravitz et al., 2011; O'Neill et al., 2012; Jégou et al., 2013).
The approach consists in analysing some key aspects of
lower-stratospheric chemistry and ozone loss in a context of high aerosol
surface area densities and high-stratospheric temperatures using
balloon-borne observations conducted in August–September 2009 from
Kiruna/Esrange in Sweden (67.5∘ N, 21.0∘ E) within the
frame of the StraPolÉté project. To our knowledge we show here the
first high-resolution in situ observations of chemical compounds obtained
within the volcanic aerosol plume of a moderate eruption. We show that in the
period on which the study is focused, N2O5 had reformed and the
role of its hydrolysis became important again after the sunlit summer period,
justifying the use of these balloon data for the investigation of
heterogeneous processes. Aerosol-constrained simulations using a 3-D
chemistry transport model (CTM) are compared to the observations. These model
calculations ignore possible dynamical effects induced by the volcanic
aerosols but are used to estimate the amplitude of the chemical impacts and
ozone loss with some comparisons with the post-Pinatubo eruption period.
Stratospheric ozone
Chemical ozone change
It is interesting to estimate the stratospheric ozone depletion induced by
the Sarychev eruption. As said above, the model does not directly calculate
possible effects of aerosols on stratospheric temperature and circulation.
All our simulations use the same transport calculations, whereas ozone loss
from Pinatubo in the northern midlatitudes can be both attributed to
chemical and transport (such as increased tropical upwelling) effects (e.g.
Telford et al., 2009; Dhomse et al., 2015). In the following, we therefore
solely calculate the change in ozone due to photochemistry.
We then compare model simulations with enhanced and background aerosol
levels (Fig. 11). Results indicate chemical reductions in ozone of a few
percent following the eruption when aerosol levels are computed from the
OSIRIS space-borne data. Accumulated ozone depletion reaches its maximum
above Kiruna near 16 km from around mid-September with changes of -22 ppbv
corresponding to -1.5 %. Below this level, changes range from -10 to
-18 ppbv, i.e. -2.5 to -3.5 %. From the upper bound of the Bal-sim
outputs, calculated ozone depletion reaches -25 ppbv (-2.8 %) and -35 ppbv
(-4 %) at 16.5 and 14 km respectively (not shown).
Changes in ozone over Kiruna (67.5∘ N, 21.0∘ E)
as a function of altitude and time between 1 July and 1 October 2009.
Calculations are done by subtracting outputs from the volcanic simulation
driven by OSIRIS observations with the background simulation. The position
of the tropopause is given by the black dotted line.
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We note that for the post-Pinatubo eruption period, ozone reductions as
large as -30 % were measured for the 12 and 22 km altitude range monitored
at some midlatitude locations in winter and spring (Hofmann et al., 1994), but these losses are both due to dynamical and chemical perturbations. Through
2-D modelling, ozone losses of up to -20 % directly resulting from
heterogeneous chemical processes were calculated in the Northern Hemisphere
lower stratosphere over the first year following the Pinatubo eruption
(Pitari and Rizi, 1993; Tie et al., 1994). The calculated chemical loss had
reduced to values much closer to those simulated for the Sarychev aerosols,
i.e. ∼ -5 %, at 60∘ N in the autumn 1992
extra-polar vortex conditions (Tie et al., 1994).
Chemical mechanisms for the ozone change in the lower
stratosphere
In the lower stratosphere, ozone removal rates are mainly controlled by the
HOx and halogen catalytic cycles which have been found to typically
account for 30–50 % and 30 % of the total ozone loss respectively in
non-volcanic conditions (Portmann et al., 1999; Salawitch et al., 2005). The
NOx cycles play a relatively minor role in the direct removal of ozone
in the lower stratosphere, but, as a result of the coupling among the
NOx, HOx and halogen cycles, the rate of ozone removal is still
very sensitive to the concentration of NOx (Wennberg et al., 1994;
Gao et al., 1999; Portmann et al., 1999; Salawitch et al., 2005). Through
the reaction of HO2 with NO (HO2+ NO → NO2+ OH),
the decreased NOx concentrations after the Sarychev eruption
result in a larger HO2 / OH ratio (as shown in Table 1) than for
background conditions (HO2 / OH ratios typically ranging from 4 to 7).
Because the photochemical removal of ozone in the lower stratosphere is
dominated by processes involving HO2, catalytic ozone destruction by
HOx cycles is likely to be amplified after volcanic eruptions (Wennberg
et al., 1994, 1995) though ozone loss rates are limited due to the
saturation of the NOx / NOy response. After the eruption of
Sarychev, the effectiveness of halogen cycles is enhanced due to increased ClOx
resulting from an OH increase (Table 1) (as explained in Sect. 4.1). However,
as said above, heterogeneous reactions activating chlorine are strongly and
non-linearly dependent on temperature, implying slow rates in the average
midlatitude temperature conditions (minimum values of 215 K) (Hanson et
al., 1994; Webster et al., 1998; Michelsen et al., 1999). Under these
conditions the simulated depletion in ozone is restrained similarly to the
finding of Tie et al. (1994) for the post-Pinatubo eruption period.
Part of the ozone depletion can be related to the coupled
BrOx / ClOx cycle, which is expected to be responsible for 20–25 %
of the halogen-controlled loss under non-volcanic aerosol conditions
(Portmann et al., 1999; Salawitch et al., 2005). Table 1 shows that the
hydrolysis of BrONO2 accounts for more than 22 % of the
ozone loss at 16.5 km after the Sarychev eruption. As described in
Sect. 4.2.2, this is due to Reaction (R4) acting as a source of OH, reducing the HCl
lifetime and thereby indirectly amplifying the chlorine-mediated ozone
depletion. Because the sticking coefficient for hydrolysis of BrONO2 on
sulfate aerosols is not temperature dependent, this effect occurs at all
latitudes and seasons in the lower stratosphere during high aerosol loading
periods (Lary et al., 1996; Tie and Brasseur, 1996).
Summary and conclusions
Our study provides key observations of the chemical perturbation in the
lower stratosphere by the moderate Sarychev volcano eruption in June 2009.
Three- and one-dimensional CTM simulations are performed to interpret balloon-borne
observations of some key chemical species made in the summer high-latitude
lower stratosphere. The modelled chemical response to the volcanic aerosols
is treated by comparing simulations using background aerosol levels and
simulations driven by volcanic aerosol amounts inferred from balloon-borne
and space-borne observations.
Quantifying the impact of volcanic aerosols on stratospheric ozone chemistry
is difficult as chemical and dynamical (radiative) effects occur simultaneously (Pitari and Rizi, 1993; Robock, 2000; Al-Saadi et al., 2001; Aquila et
al., 2013). The model is a CTM driven by ECMWF offline meteorological data
and does not describe radiative processes. In other words, volcanic aerosol
radiative effects do not directly interact with the circulation computed
by the model. Radiative processes from the injection of volcanic aerosols
into the tropics have been shown to have an impact on mean meridional circulation
and ozone transport (Brasseur and Granier, 1992; Pitari et Rizi, 1993). In
our study, effects of the Sarychev aerosols on midlatitude stratospheric
dynamics, if any, are at least of the first order intrinsically taken into
account in the ECMWF analyses used for all simulations. REPROBUS does not
take into account the aerosol impact on calculated photolysis rates, which is
likely to result in some differences between models when this process is
computed or ignored (Pitari and Rizi, 1993; Pitari et al., 2014). However,
because the Sarychev eruption has only impacted the lower stratosphere at
mid- and high latitudes the effect on the photolysis frequency of molecular
oxygen and ozone due to absorption and backscattering of solar radiation by
the volcanic aerosols is expected to be very small in these regions (Tie et
al., 1994). Therefore, since all our simulations were driven with the
same wind and temperature fields, our approach only estimates the chemical
effects of the Sarychev aerosols.
The NOy chemistry appears to be very sensitive to the increase in SAD
within the lower stratosphere resulting from the Sarychev eruption. A
decrease in the NOx abundances is evident but shows some saturation as
emphasized in a number of studies referring to cases of high sulfate aerosol
loadings (e.g. Fahey et al., 1993). The effect of volcanic aerosols on
nitrogen partitioning is also reflected in the calculated production of
HNO3 as a result of the decrease in the N2O5 nitrogen
reservoir from its enhanced hydrolysis and NOx reduction.
Although direct comparisons in terms of solar illumination, latitude,
injection altitudes and temperature are not possible for distinct volcanic
eruptions such as Pinatubo and Sarychev, it is interesting to compare the
effect of both eruptions on the photochemistry of the lower stratosphere.
Overall, although different in magnitude, the eruptions of Pinatubo and
Sarychev show a similar observed and simulated depletion of NO2, probably
due to the saturation effect of the enhanced N2O5 hydrolysis. In
comparison with the Pinatubo period, the Sarychev aerosols led to less
overall HNO3 production in the stratosphere, possibly because the
related HNO3 enhancement has been shown to be considerably weaker in
the lowermost stratosphere (below ∼ 18 km) than for sulfur
injection into higher altitudes (Webster et al., 1994; Santee et al., 2004).
However, one should note that previously reported modelling studies on the
Pinatubo aerosols were conducted with previous chemical kinetic rate constants
and photolysis rates which have mostly been updated since, which adds a degree of complexity to comparisons discussed within the present study.
For the Pinatubo aerosols, ozone destruction was not observed throughout the
volcanic aerosol layer because N2O5 hydrolysis reduced NOx-related ozone loss, which even resulted in small increases in ozone in the
middle stratosphere (Bekki and Pyle, 1994; Tie and Brasseur, 1995). For the
Sarychev eruption, the volcanic aerosol layer is restrained to altitude
levels below 19 km where the ozone destruction processes by HOx and
halogen catalytic cycles are expected to play a major role (e.g. Salawitch
et al., 2005), with some sensitivity towards NOx levels. To summarize,
the increased production of HNO3 via N2O5 hydrolysis enhances
the photolytic production of OH from HNO3. As a result, the gas-phase
sink for HCl by reaction with OH is slightly enhanced and is associated with
an increase in ClO amounts. An important result of the heterogeneous
hydrolysis of BrONO2 is the formation and subsequent photolysis of
additional HOBr. The OH so produced additionally converts HCl to ClO (and
ultimately to ClONO2). Accordingly, there is substantial repartitioning
of the active chlorine, but effects of the BrONO2 hydrolysis on
nitrogen partitioning are insignificant. In this chemical context, the
magnitude of the ozone response to the Sarychev volcanic perturbation
appears restricted (for instance, -22 ppbv or -1.5 % around 16 km) because
the saturation of the NOx / NOy response limits the increase in
HOx and in active chlorine (ClO) by enhanced HOx, precluding
important ozone loss rates. Moreover, stratospheric temperatures remained
too high (i.e. mainly above 215 K) for efficient heterogeneous conversion of
ClONO2 to active chlorine, which could have led to significant ozone
depletion. For these temperature conditions, Reaction (R2) is not expected to
compete with N2O5 hydrolysis in the NOy partitioning (Fahey
et al., 1993; Cox et al., 1994).
However, limitations in our model simulations also contribute to some
model–measurement discrepancies. A first major difficulty is to drive the
model simulations with representative and consistent inputs in term of
volcanic aerosol loading. To address this issue, two different model runs
for aerosol forcing have been performed, one using OSIRIS satellite data
converted to aerosol SAD fields and the other one from in situ balloon-borne
observations. The OSIRIS satellite data represent zonally and daily averaged
values of SAD, which may vary from a 3-D construction based on the local
surface areas. The possible presence of aerosol streamers (geographical
variations in the aerosol content) resulting from the transport of the
volcanic aerosols over the Northern Hemisphere present from mid-July to
September 2009 is likely to affect the N2O5
abundances locally and regionally and, to a lesser extent, NO2 and HNO3 (Jucks et al.
1999; Küll et al., 2002). If our aerosol SAD dataset had been obtained
when the local concentrations were higher than the zonal mean values, then
the calculated rate of the heterogeneous reactions would be biased low and
calculated NOx and HNO3 abundances would be systematically biased
high and low respectively. This is not, however, evident in all our
comparisons from simulations based on OSIRIS aerosols. The second type of
aerosol-constrained simulation uses SADs from balloon-borne observed
profiles. By definition, such in situ observations deal with a particular
location. Extrapolating in situ derived SADs to drive a 3-D model on a large
scale may induce inaccurate simulations of the chemical impact of the
aerosols (Kondo et al., 2000). To account for this SAD-related uncertainty,
our simulations based on in situ data encompass the range of SADs derived
from the STAC balloon-borne observations over the August–September 2009
period. Both satellite- and balloon-driven simulations give similar results
in terms of NO2 and HNO3 amounts, possibly because the in situ
observations represent the aerosol loading at the northern
midlatitudes well. Another explanation is that the saturation effect (roughly
when SADs become larger than 3 µm2 cm-3) of the
NOx / NOy ratio is more relevant for the range of observed SADs than
spatiotemporal inhomogeneities.
Secondly, adequate modelling of transport is also crucial for the
partitioning of NOy. Processes that control the vertical profiles of
NO2 and HNO3 in the stratosphere are based on a complex interplay
between dynamics and chemistry with the key issue being to accurately simulate
total NOy, which may be not systematically achieved with 3-D CTM
calculations. Improved simulations of transport can be obtained by combing
operational analyses with forecasts to construct 3-hourly meteorological
data to drive the CTM (Berthet et al., 2006). We have applied this strategy
in the present study. Using 1-D modelling driven by in situ observations or
calculating NO2 / HNO3 ratios to reduce transport effects does not
clearly improve the model–measurement comparisons for the lower
stratosphere. Although some features in the vertical profiles are not
systematically captured by the model, this tends to indicate that the error
in calculated transport is not large enough to account for the overall
difference between measured and modelled NO2 and HNO3 when no
volcanic aerosol loading is included in the model. Rather, these results
show some evidence of the role of heterogeneous reactions at the surface of
volcanic aerosols.
Thirdly, part of the discrepancies between model and observations might be
attributed to spatial resolution issues. It may be tricky to compare model
calculations with high-resolution in situ profiles and with remote-sensing
observations integrating over tens of kilometres (Berthet et al., 2007). For
instance, discrepancies between remote-sensing observations and model
calculations have been reported for stratospheric NO3 in the case of
localized temperature inhomogeneities as a result of the strong dependence
of NO3 cross sections and kinetics on temperature (Renard et al.,
2001). N2O5 and NO2 may be subsequently impacted because
NO3, together with NO2, plays a central role in the equilibrium
reaction controlling N2O5 in the gas phase.
In our study, no comprehensive sulfur chemistry is included in the model. We
have also excluded dynamical and radiative effects on the ozone response, which have been shown to be of primary importance when dense volcanic clouds
are present (e.g. Pitari and Rizi, 1993; Kinnison et al., 1994; Tie et al.,
1994; Al-Saadi et al., 2001). In a forthcoming study it would be interesting
to compare dynamical or radiative and chemical effects of moderate volcanic
eruptions on stratospheric ozone using chemistry–climate models with full
sulfur chemistry and aerosol–dynamics interactive calculations.
Finally, it might be interesting to investigate the effects of other
volcanic plumes coming from moderate volcanic eruptions which are then
transported to high-latitude regions when stratospheric temperatures are
more favourable for chlorine activation and enhanced ozone loss (e.g. in
winter). The activation of chlorine from volcanic sulfate aerosols and
associated ozone depletion is arguably more significant in the cold
temperature conditions of winter and spring, even above the formation threshold
of polar stratospheric clouds (Hanson et al., 1994). The eruption of the
Calbuco volcano in the Southern Hemisphere in April 2015 could be a good
candidate for a study of this process (Solomon et al., 2016).