Introduction
Energetic particles enter the Earth's atmosphere near the magnetic poles
altering the chemistry of the middle and upper atmosphere. Energetic particle
precipitation (EPP) is the major source of nitrogen oxides (NOx) and
hydrogen oxides (HOx) in the polar middle and upper atmosphere
. Both chemical components catalytically
deplete ozone; NOx mainly below 45 km and HOx mainly above.
HOx is short-lived in the middle atmosphere and depletes ozone mainly in
the mesosphere. In contrast, NOx persists for up to several months in the
polar winter middle atmosphere. Inside the polar vortex, NOx can be
transported downward from the lower thermosphere to the stratosphere, where
it depletes ozone e.g.,.
Observational evidence of polar winter stratospheric ozone loss due to EPP is
still limited. Only recently have long-term satellite observations with good
temporal and spatial coverage become available. In austral polar winter EPP
causes an ozone loss of about 10–15 % descending from 1 hPa in early
winter to 10 hPa in late winter .
Extensive information on the current knowledge of energetic particle
precipitation can be found in and
.
Ozone loss influences stratospheric temperature and the polar vortex. The
Northern Annular Mode (NAM) index is often used to describe the strength of
the polar vortex, with positive NAM values indicating a strong polar vortex
and negative NAM values indicating a weak polar vortex. Observations indicate
that anomalous weather regimes associated with the NAM index can propagate
from the stratosphere down to the surface . Hence,
energetic particle precipitation may provide a link from space weather to
surface climate. Here, we study the impact of ozone loss due to EPP on the
circulation and subsequently on climate. Discussed are both a polar
mesospheric and a polar stratospheric ozone loss.
Since the discovery of the ozone hole in the mid-1980s, the climate impact of
stratospheric ozone loss has been intensively studied
e.g.,. Most studies
concentrated on the climate impact of the ozone hole during austral spring
and have reported cooling in the spring Southern Hemispheric stratosphere due to
reduced absorption of solar radiation and strengthening of the polar
vortex. In contrast, our study concentrates on ozone loss during the
boreal polar night. During polar night, reduced ozone slightly decreases the
infrared cooling of the polar stratosphere resulting in net (small)
stratospheric warming . However, both studies
prescribed ozone loss in the lower stratosphere.
Several studies have suggested a significant influence of EPP on climate.
and used reanalysis data to investigate
the dependence of stratospheric temperature and zonal wind on the Ap index.
They found stratospheric warming of up to 5–10 K for strong energetic
particle precipitation descending from the stratopause to the
mid-stratosphere. However, for the zonal wind response the two studies differ
from each other. found strengthening of the polar
vortex, whereas showed weakening of the polar vortex.
Moreover, analyzed surface air temperature changes in
reanalysis data for years with various strengths of EPP. They found warming
over Eurasia and cooling over Greenland for winters with enhanced EPP, but
could not rule out that the estimated changes are induced by NAM variability
independent of EPP.
Other studies relied on atmospheric chemistry models, which showed similar
surface temperature change patterns as found in the reanalysis data
e.g.,. They
reported small cooling in the polar winter stratosphere due to EPP.
However, the radiative effect of a polar night ozone loss should lead to warming, which can also be found in reanalysis data
. The simulated stratospheric cooling is
attributed to dynamical, adiabatic cooling caused by a decrease in the mean
meridional circulation .
suggested that the weaker mean meridional circulation
is caused by a decrease in midlatitude tropospheric wave forcing. The
aforementioned model studies analyzing the climate impact of EPP relied on
relatively few simulation years and applied complex forcings. Instead of
prescribing ozone, these studies simulated EPP effects by changing the
production of NOx and HOx and modeling the effects on ozone
interactively. This could potentially be more realistic than simulations with
prescribed ozone anomalies, but it introduces uncertainties related to the
representation of chemistry and transport in the model, and renders the
understanding of the effects more complicated as the ozone forcing varies in
space and time. To avoid these difficulties and to obtain a clear
signal-to-noise ratio, we use an idealized ozone forcing and a long
simulation period.
Commonly, the effects of EPP are classified into direct and indirect effects
. Direct effects are those of the
local production of NOx and HOx, whereas indirect effects are the
effects of the NOx transport from the thermosphere to the stratosphere.
Whereas most of the abovementioned studies discuss a mainly stratospheric
ozone loss due to the indirect EPP effect, suggested a
potential climate influence of mesospheric ozone loss due to the direct EPP
effect. By using satellite observations they showed that HOx causes
long-term variability in mesospheric ozone of up to 34 % between EPP maximum
and EPP minimum. were the first to include the direct
effect of HOx local production due to EPP in a chemistry–climate model.
They found a similar mesospheric ozone loss as and
ultimately, reported cooling over Greenland and warming over Eurasia.
However, also considered the indirect effect of the
NOx descent. Hence, the sole impact of mesospheric ozone loss due to the
direct EPP effect as suggested by remains unclear.
This paper studies the circulation and climate impact of idealized
mesospheric and stratospheric ozone losses that could be attributed to
energetic particle precipitation. We use simulations with the Max Planck
Institute Earth System Model (MPI-ESM), applying an idealized ozone forcing in
either the mesosphere or the stratosphere. The idealized mesospheric ozone
loss that we prescribe may be considered to be mostly a direct EPP effect,
whereas the prescribed stratospheric ozone loss should be considered
indirect. Additionally, we use a radiative transfer model to quantify the
radiative forcing of ozone at different altitudes and months. Ultimately, we
discuss whether ozone loss in the middle atmosphere due to EPP has the
potential to significantly alter the surface climate.
Section describes the MPI-ESM as well as the
radiative transfer model. Section links mesospheric
and stratospheric ozone losses to changes in the atmospheric temperatures and
winds. Finally, Sect. summarizes and discusses
the main outcomes and limitations of this study.
Models and numerical experiments
MPI-ESM: the Max Planck Institute Earth System Model
The MPI-ESM consists of the coupled atmospheric and ocean
general circulation models, ECHAM6 and MPIOM
, the land and vegetation model JSBACH
, and the model for marine bio-geochemistry HAMOCC
. We use the “mixed-resolution” configuration of the
model (MPI-ESM-MR). The ocean model uses a tripolar quasi-isotropic grid with
a nominal resolution of 0.4∘ and 40 vertical layers. ECHAM6 is run
with a triangular truncation at wave number 63 (T63), which corresponds to
1.9∘ in latitude and longitude. The vertical grid contains 95 hybrid
σ-pressure levels resolving the atmosphere from the surface up to
0.01 hPa. The vertical resolution is nearly constant (700 m) from the upper
troposphere to the middle stratosphere, and is less than 1000 m at the
stratopause. The time steps in the atmosphere and ocean are 450 and 3600 s,
respectively.
The model has been used for many simulations within the CMIP5 (Coupled Model
Intercomparison Project Phase 5) framework . An overview
of the dynamics of the middle atmosphere in these simulations is given by
. In this study, the preindustrial CMIP5 simulation
(piControl) is used as reference. The forcing is constant in time and uses
pre-industrial conditions (AD 1850) for the greenhouse gases. Solar
irradiance and ozone concentrations are averaged over a solar cycle
(1844–1856 for the solar irradiance and 1850–1860 for ozone
concentrations). No volcanic forcing is applied. A period of 150 years of
this simulation is used.
In order to analyze the impact of ozone changes on the model climate, two
additional experiments with reduced ozone concentration are carried out. In
one experiment, the mesospheric ozone is reduced by 40 % between 0.01 and
0.1 hPa polewards of 60∘ N (this is called “meso-O3”). In the
other experiment, stratospheric ozone is reduced by 20 % between 1 and
10 hPa polewards of 60∘ N (this is called “strato-O3”). We
perform on–off experiments, whereas in reality EPP causes a constant (but
variable) ozone loss. However, the magnitude of the prescribed ozone losses
is based on satellite observations for winter conditions between years with
high geomagnetic activity and years with low geomagnetic activity. In
general, the impact of energetic particles is sporadic in the mesosphere;
, however, showed that the direct HOx effect induces
long-term variability in mesospheric ozone of up to 34 % from November to
February in satellite data. and
revealed an upper stratospheric ozone loss between 10 and 15 % due to
energetic particles for the Antarctic high latitudes in long-term
measurements. Note that the applied ozone losses are slightly larger than the
EPP influence diagnosed from observations. We use the stronger forcing to
obtain a clear signal-to-noise ratio. However, this implies a potentially
overestimated climate response.
To facilitate experiment design, we fixed ozone losses to be constant
over time. Although we concentrate our analysis on boreal winter high
latitudes, this still allows us to gain insights into boreal spring (i.e., the
transition time from polar night to polar day). Observed ozone losses in
summer are in general smaller than during winter, but this idealized setting
allows for easy comparison of potential effects during the different seasons.
In order to test whether the winter response is
influenced through preconditioning, in this experiment design we repeated the experiments with ozone
losses prescribed only from December to March. However, as the results are
qualitatively very similar and differ only in the magnitude of the responses,
we discuss only the results of the experiments with ozone losses prescribed
all year. Both experiments, with mesospheric and stratospheric ozone loss,
are forced by the same conditions as the piControl experiment. Moreover, the
simulations are restarted from the same year in the piControl experiment.
This ensures that the ocean state is similar in all experiments. For both
simulations 150 years are simulated.
The simplistic nature of our experiments is intentional and, we think, useful.
We chose this idealized experimental design in order to separate the climate
impact of stratospheric and mesospheric ozone loss due to EPP and to identify
the relevant mechanism which determines how EPP affects the climate. Prescribing complex ozone
reductions that vary in space, by season and year, or
simulating the ozone reduction interactively might enable more realism, but
they do not facilitate the identification of potential mechanisms. However, due to
the simplification we cannot consider all features associated with EPP. In
particular, three main effects are not taken into account: (a) energetic
particles entering the atmosphere only over the auroral oval regions
; (b) the negative ozone signal due to
EPP propagating from the stratopause in mid-winter to the lower stratosphere
in spring within the polar vortex ; and
(c) that the polar vortex can shift off the pole to regions with more solar
radiation. We, instead, apply constant ozone reduction between the
stratopause and mid-stratosphere (1–10 hPa) over the entire polar cap. The
climate response in our simulations is likely overestimated as we reduce
ozone over a larger latitudinal and altitude region than observations
suggest.
In Sects. and the
differences between the experiments and the control simulation (i.e.,
piControl) are analyzed. Statistical significance is calculated using the
95 % confidence intervals assuming normally distributed regression errors
and using the 0.975 and 0.025 percentile of Student's t distribution with
the appropriate degrees of freedom. Properties of two simulations are
considered statistically significantly different if the mean value of the
control simulation is outside 95 % confidence interval of the experiment.
The radiative transfer model PSrad
The radiative transfer scheme of MPI-ESM is based on the rapid radiation
transfer suite of models optimized for general circulation models (RRTMG;
). The RRTMG is widely used and its ability
to calculate radiative forcing has been evaluated by . In
its stand-alone version here, it is used to study the impact of ozone on
heating rates. It is divided into 16 bands in the longwave
(1000–3 µm) and 14 bands in the shortwave (12 195–200 nm; ). The spectral bands are chosen to include the major
absorption bands of active gases. The major ozone absorption bands – Hartley
band (200–310 nm), Huggins bands (310–350 nm), and Chappuis bands
(410–750 nm) – are considered. However, absorption of oxygen at shorter
wavelengths than 200 nm is missing, which could lead to an underestimation
of the total heating rate in the mesosphere. The radiative transfer scheme is
further described in and , and
further on we will refer to it as the radiative transfer model “PSrad”.
The shortwave and longwave components are calculated separately. Furthermore,
optical properties for gases, clouds, and aerosols are computed separately for
longwave and shortwave components and, finally, combined to compute the total heating
rates. PSrad expects profiles of gases (H2O, N2O, CH4, CO, O3),
profile of cloud parameters as well as additional parameters (e.g., albedo
and zenith angle) as input. Additionally, CO2 and O2 are set to fixed
values invariant with height. For all other gases, we use multi-year monthly
means representative of the late 20th century provided by the atmospheric
and chemistry model HAMMONIA (Hamburg Model for Neutral and Ionized
Atmosphere; ). For the albedo and cloud properties
(e.g., cloud fraction, cloud water/ice content), multi-year monthly means
from the piControl experiment are used. All quantities are extracted for
75∘ N. The zenith angle is calculated for 12:00 UTC at
75∘ N, 0∘ E for the 15th of each month. The latitude of
75∘ N is chosen to be exemplary for polar latitudes. The results are
insensitive to the actual latitude, the main difference at other polar
latitudes is the length of the polar night. Note that the length of the polar
night for an air pocket also depends on the altitude and on atmospheric
dynamics (e.g., movement of the polar vortex). Both effects are omitted in
this study. In our simulations ozone is reduced independent of actual
dynamics, over the entire polar cap (60–90∘).
To quantify the impact of ozone on the heating rates, we perform multiple
runs in which for each run the ozone concentration of a single layer is set
to 0. Then we take the differences between a control run and each single
run. The differences of each run are, finally, added up for the estimation of
the total heating rate. This method allows us to consider that layers of
reduced ozone will lead to increased absorption of shortwave radiation in the
layers directly below.
Results
Ozone effects on the heating rates
Ozone loss directly alters the atmospheric energy transfer. Before
analyzing circulation and climate impacts due to ozone losses, we study the
heating rate response using the radiative transfer model PSrad. The heating
rates are calculated for the polar latitude of 75∘ N (see
Fig. ). As the effect of EPP is most important at the
winter polar cap, we will concentrate our analysis on boreal winter high
latitudes.
In the shortwave part of the spectrum, ozone strongly absorbs solar radiation
and heats the whole atmosphere. The strongest heating (about
12 K day-1) occurs in the uppermost stratosphere around 1 hPa. Ozone loss would, hence, result in relative cooling due to reduced heating.
The ozone heating and, hence, the cooling caused by ozone reduction become smaller for larger zenith angles and vanish in polar night.
In the longwave part of the spectrum, the radiative effect of ozone is highly
temperature dependent. Ozone cools the atmosphere via infrared emission in
the stratosphere and in warm regions of the mesosphere below 0.1 hPa (see
Fig. b). The strongest cooling (about
-2 K day-1) occurs at the stratopause. In the troposphere and in the
cold regions of the mesosphere above 0.1 hPa, the absorption of outgoing
radiation exceeds the infrared emission resulting in a heating of the
atmosphere due to ozone.
Monthly mean heating rates of ozone (K day-1) for
75∘ N calculated by the radiative transfer model PSrad
for (a) shortwave, (b) longwave, and (c) total
(shortwave + longwave) radiation.
In total, the shortwave heating dominates all sunlit months. During polar
night, ozone cools the atmosphere between 0.1 and 100 hPa and, hence, ozone loss in the stratosphere and lower mesosphere results in warming.
Near the terminator (e.g., at 75∘ N in November and February), the
net influence of ozone is more complex: at some altitudes ozone heats the atmosphere and at
some ozone cools it. The net radiative forcing of ozone loss
depends on when and where ozone is reduced. For example, in November, stratospheric (1 hPa) ozone loss leads to a heating, but mesospheric
(0.1 hPa) ozone loss leads to cooling.
These results are in line with previous work. It is widely accepted that ozone loss in spring and summer leads to stratospheric cooling
e.g.,. Some studies analyzed the radiative
forcing of winter stratospheric ozone loss. showed that
the observed stratospheric ozone loss in the late 20th century led to winter warming and summer cooling
in a global climate model. Using a radiative transfer
model with fixed dynamical heating, confirmed that stratospheric ozone loss over the winter pole results in small
stratospheric radiative warming and dominating stratospheric dynamical
cooling. showed that the shortwave cooling of the
stratosphere due to ozone loss dominates, in all sunlit months, infrared
heating due to ozone loss. Recently, simulated warming in mid-winter and cooling in late winter and spring in the upper
stratosphere for ozone losses explicitly induced by EPP.
The results above are confirmed by the actual heating rate anomalies
induced by the applied ozone losses in the experiments meso-O3 and
strato-O3 (not shown). The heating rates are calculated at the first
time step of the model at which the radiation is updated (1 January)
excluding any feedbacks occurring only at later time steps. Note that the
exact values may change for other months, especially depending on the sunlit
area. Compared to the total heating rates of piControl, the changes in
heating rates caused by a 40 % reduction of mesospheric ozone in polar
night are very small (on average about -0.01 K day-1 and
-0.4 %),
and by a 20 % reduction of stratospheric ozone on average are about
0.12 K day-1 and 2.6 %. This agrees with the estimate of
, who simulated a change of 0.1 K day-1 in the
winter stratospheric heating rate due particle-induced ozone loss. The change
in heating rates due to the stratospheric ozone change is in the range of
solar UV forcing (0.1 K day-1) .
Zonal mean temperature (K; top row) and zonal mean
zonal wind (m s-1; bottom row) averaged over December–February (DJF)
for (a, d) piControl, (b, e) the difference between
meso-O3 and piControl, and (c, f) the difference between
strato-O3 and piControl. Shaded areas are statistically significant at the
95 % confidence interval. The black, dashed boxes highlight the regions
where ozone is reduced.
Climate effects of mesospheric ozone loss
As changes in heating rates due to reduced ozone during polar night are
small, one might reason that climate impact of winter polar ozone loss is
small. But large effects may occur in regions slightly outside the polar
night. Furthermore, several studies suggested that changes in the heating
rates due to winter polar ozone loss leads to dynamical cooling
e.g.,, whereas the
initial radiative forcing suggests warming. Hence, we further analyze the
climate impact of winter polar ozone loss. As large variations in the polar
vortex can propagate downward and affect the surface climate, we first
concentrate on the circulation changes of the middle atmosphere due to ozone loss, which are a prerequisite for a potential climate impact of EPP.
In the following, we analyze the climate effect of idealized polar
mesospheric ozone loss, while in Sect. we analyze
the climate effect of idealized polar stratospheric ozone loss.
Figure a and d show the zonal mean
temperature and zonal wind simulated for boreal winter (December–February).
Main observed characteristics of the zonal mean temperature, e.g., the
stratopause tilt from the summer towards the winter pole, are well
reproduced. The changes in the zonal mean zonal wind are consistent with the
temperature changes via the thermal wind balance. In most regions, the
difference between meso-O3 and piControl is very small (see
Fig. b and e).
Near the winter pole, a dipole structure emerges with cooling in the upper
stratosphere and warming in the mesosphere. According to our radiative
transfer calculations mesospheric winter polar ozone loss should lead to cooling. However, the temperature differences are small (below 1 K) and not
significant at the 95 % level. As the applied forcing is very small, small
and barely significant values are expected. At the winter pole, the polar vortex
slightly weakens, whereas the mesospheric winds strengthen; these differences
are not significant. The signal is only slightly stronger but still
insignificant if winters with major sudden stratospheric warmings (SSWs) are
excluded (not shown). As stated above, large variations in the winter polar
vortex can propagate to the surface influencing the surface climate. However,
the changes reported here are small. The anomalies reaching the troposphere
are statistical artifacts. Indeed, the surface temperature reveals no
statistically significant change (not shown).
The temperature and wind signals are not statistically significant
after 150 simulated years; it nevertheless makes sense to analyze if the
signals could have a physical explanation and not be purely accidental. Note
that with fewer simulation years apparently very different results can be
obtained. Upon analyzing different simulation periods, we obtain mesospheric
warming and cooling of apparent significance. Particularly, we calculated a
statistically significant weakening of the polar vortex when using only the
first 80 simulation years. We cannot identify a model drift in the
experiments, which could explain the disagreement between the 150-year and
80-year runs. However, the model simulates variability on timescales up to
multi-decadal, which is common in many climate models ,
and might cause the apparently different responses to ozone reduction in
different sub-periods of the 150-year simulation. The high degree of internal
variability of the winter polar stratosphere can obviously create incorrect
apparent signals. The most dramatic demonstration of this variability are
major sudden stratospheric warmings (SSWs), which occur on average about six
times per decade in the Northern Hemisphere (see , for
more information on SSW). A short simulation period may lead to an
over-representation or under-representation of SSWs. Over our whole
simulation period (150 years) the number of major SSWs is balanced in all
three experiments. In total, there are 102 events in piControl, 99 events in
meso-O3, and 109 events in strato-O3 (using a reversal of the zonal wind
at 60∘ N and 10 hPa as the criterion for major SSW occurrence).
Climate effects of stratospheric ozone loss
In this section, we analyze the climate effect of idealized polar
stratospheric ozone loss. Figure a and d
show the zonal mean temperature and zonal wind simulated for boreal winter
(December–February) for piControl, and Fig. c and
f show the difference between strato-O3 and piControl. The
winter stratosphere warms due to ozone loss as expected from the
calculations with the radiative transfer model. As a consequence of the
warming, the stratospheric winds weaken. The small mesospheric cooling likely
results from enhanced eastward momentum deposition from gravity waves as
shown by . Our results are in line with earlier studies.
and identified warming in the polar
winter upper stratosphere due to EPP in reanalysis data, but their magnitude
is much stronger (5 K) than in our simulations. Regarding the zonal wind
response, the two studies differ from each other.
analyzed strengthening of the polar vortex with enhanced equatorward
planetary waves, whereas analyzed weakening of the polar
vortex. The statistically significant warming of the summer mesopause is an
indication of inter-hemispheric coupling as discussed by
and also persists for winters without a sudden
stratospheric warming event.
Monthly mean temperature (top row) averaged between 60 and
90∘ N (K) and zonal wind (m s-1; bottom row) for
60∘ N for (a, c) piControl and (b, d) the
difference between strato-O3 and piControl. Shaded areas are statistically
significant at the 95 % confidence interval.
Surface temperature (K) averaged over December–February (DJF) for
the Northern Hemisphere for (a) piControl and (b) the
difference between strato-O3 and piControl. Shaded areas are statistically
significant at the 95 % confidence interval.
Figure shows only changes for the mean over December to
February, while the radiative transfer model suggests that the month-to-month
variability of the forcing is large. To study whether the impact of stratospheric ozone loss differs over the course of the winter, we analyze
the monthly means of temperature and zonal wind (see
Fig. ). Ozone loss during most of the polar night
(except December) leads to warming, whereas at all other times and
locations it leads to cooling. This agrees with the calculations of the
radiative transfer model and with our assumption that the winter cooling is
not affected by strong summer warming. However, the cooling in December is
unexpected from the radiative transfer modeling. argued
that the polar winter atmosphere transits from a radiatively controlled state
in early winter to a dynamically controlled state in late winter. Given the
opposite sign of the diabatic forcing, the simulated cooling must have been
already dynamically created in December. This is in agreement with early model
studies which showed that uniform ozone losses lead to dynamical cooling at
the boreal winter polar latitudes e.g.,.
suggested that the dynamical cooling is due to weakening of the mean meridional circulation related to reduced wave forcing
caused by a reduction of midlatitude wave flux into the stratosphere.
Similarly, in our simulations we find a (albeit not significant) reduction of
the zonal mean eddy heat flux at 100 hPa in the midlatitudes from December
to March (not shown). This may be caused by enhanced wave reflection as
suggested by for the dynamical response to 11-year solar
irradiance forcing. The dynamically induced cooling in December also occurs
in simulations in which the ozone is only reduced from December to March (not
shown). Also, reported dynamical cooling in the
winter polar stratosphere due to EPP. However, in their model the cooling
dominates the winter (DJF) signal, whereas we obtain small warming for the
DJF average (see Fig. ). The magnitude of the signal
decreases in our simulations, especially in late winter, if we exclude all
seasons with a SSW (not shown).
The zonal wind changes consistently with the temperature changes via the
thermal wind balance. Simultaneously with warming (cooling), the polar
wind weakens (strengthens). Anomalies in the polar vortex occasionally reach
the troposphere (e.g., the strengthening in November or the weakening in
December or February). Although most of those changes are not significant,
some disturbances in the polar vortex may still force the surface temperature
(see Fig. ). In our simulations for boreal winter,
stratospheric ozone loss cools large parts of the northern high latitudes
from northern Europe to Eurasia and over North America. Excluding all
winters with a SSW strengthens the cooling over North America (not shown).
Over Greenland and the pole, the surface warms. This is consistent with the
weakening of the zonal wind in December (see
Fig. d). However, most changes are small and not
significant. and analyzed
statistically significant changes in surface temperature and found warming over
Eurasia of about 1.5 K and cooling over North America of about -1 K.
Compared to both studies the amplitude of our signal is much smaller. The
weaker signal also persists if we exclude all winters with a SSW (not shown).
However, based their study on only 9 simulated
years and we have shown that the large variability in the polar winter
stratosphere can cause incorrect apparent signals if the ensemble is not large
enough. Note that another difference between our and their study is that our
model is coupled to an interactive ocean. could not rule
out that their results are by chance induced by the NAM.
Summary and conclusion
In this study, we analyzed the climate impact of idealized mesospheric and
stratospheric ozone losses. Although this study is motivated by the
enhancement of NOx due to energetic particle precipitation (EPP), the
results presented here could also be applied to other processes causing ozone
destruction. We lie the focus on boreal winter. The radiative forcing of
polar ozone is calculated by the radiative transfer model PSrad. In
sensitivity studies with the Max Planck Institute Earth System Model
(MPI-ESM), we applied idealized ozone losses of either 40 % in the winter
polar mesosphere or 20 % in the winter polar stratosphere. This
simplified design facilitates the identification of the processes relevant
for possible climate responses.
Recently, showed that the direct EPP-HOx effect
induces large long-term variability in winter mesospheric ozone. They
suggested that these large changes may have an impact on climate. Following
their idea, we analyzed the atmospheric response to mesospheric ozone loss.
We found that the winter atmospheric changes due to mesospheric ozone loss
in our model are negligible. Calculations with a radiative transfer model
showed that the radiative forcing of mesospheric ozone is very small during
polar night, which makes the small dynamic response plausible.
Several studies have analyzed the climate effect of stratospheric ozone loss due
to EPP. calculated a correlation of the winter surface
temperature and energetic particle precipitation in reanalysis data. However,
they could not rule out accidental correlation. Since
then several model studies have tried to establish a physical link between EPP and
climate . In all these
model studies, dynamical cooling of the winter polar stratosphere due to
energetic particle precipitation was simulated. In our model, stratospheric
ozone loss during polar night (except December) results in warming, whereas
at all other times and locations it leads to cooling. This agrees with the
calculations of the radiative transfer model. We found cooling during
December due to stratospheric ozone loss caused by reduced vertical wind.
However, the changes in the polar winter stratosphere are small and not
significant in our model. Consequently, the impact on the simulated
winter surface temperature is also weak. In contrast to the abovementioned
studies, in our experiment the dynamical feedback leading to the
stratospheric cooling is not dominant throughout the boreal winter. This is
also true if we restrict the ozone loss to December to March. However, the
earlier model studies were based on only a few simulation years. Using only
the first 80 years of our simulations we obtained false positives. The high
degree of internal variability of the polar vortex can create incorrect apparent
signals.
As the radiative forcing of our prescribed mesospheric ozone loss is
negligible, a significant climate impact of mesospheric ozone change as
suggested by seems unlikely. Our experimental design
would likely rather overestimate the climate impact of EPP than underestimate
it. However, our simulations indicate only small changes in the stratospheric
circulation and temperature and a weak impact on surface temperature. We
encourage more research on the effects of EPP as the climate impact of
stratospheric ozone losses due to EPP is not as clear as often thought and
the underlying processes are not well understood. The upcoming CMIP6 model
intercomparison may help to resolve those open points, because energetic
particle forcing is recommended – for the first time – as part of the solar
forcing . The role of wave reflection for the
coupling mechanism between stratosphere and troposphere, in particular, needs to be
clarified. Furthermore, the catalytic destruction of ozone by NOx works
only effectively if sunlight is available. The influence of EPP-induced
NOx may be larger near the terminator.
Moreover, the simplified experimental design has its limits. It is suitable
to address the different processes related to direct and indirect EPP impacts
and the identification of mechanisms for possible climate responses. However,
we cannot rule out that the time and altitude dependence of the ozone loss
caused by the downward transport of ozone and nitrogen oxides in the polar
vortex is important, but we obtain qualitatively very similar results if the
ozone is only reduced during December to March.
Finally, although previous studies have shown that MPI-ESM reproduces
stratospheric temperature responses to forcings reasonably well
e.g.,, the possibility remains that the
model's sensitivity to ozone loss is biased low. To address this, we would
encourage multi-model studies on EPP climate impact as currently
suggested for the third phase of the SOLARIS-HEPPA project, which
investigates solar influences on climate as part of the
“Stratosphere-troposphere Processes And their Role in Climate” (SPARC)
project.