The atmospheric response to the 11-year solar cycle is separated into the
contributions from changes in direct radiative heating and photolysis rates
using specially designed sensitivity simulations with the UM-UKCA (Unified
Model coupled to the United Kingdom Chemistry and
Aerosol model) chemistry–climate model. We perform a number of idealised time-slice
experiments under perpetual solar maximum (SMAX) and minimum conditions
(SMIN), and we find that contributions from changes in direct heating and
photolysis rates are both important for determining the stratospheric
shortwave heating, temperature and ozone responses to the amplitude of the
11-year solar cycle. The combined effects of the processes are found to be
largely additive in the tropics but nonadditive in the Southern Hemisphere
(SH) high latitudes during the dynamically active season. Our results
indicate that, in contrast to the original mechanism proposed in the
literature, the solar-induced changes in the horizontal shortwave heating
rate gradients not only in autumn/early winter but throughout the
dynamically active season are important for modulating the dynamical
response to changes in solar forcing. In spring, these gradients are
strongly influenced by the shortwave heating anomalies at higher southern
latitudes, which are closely linked to the concurrent changes in ozone. In
addition, our simulations indicate differences in the winter SH dynamical
responses between the experiments. We suggest a couple of potential drivers
of the simulated differences, i.e. the role of enhanced zonally asymmetric
ozone heating brought about by the increased solar-induced ozone levels
under SMAX and/or sensitivity of the polar dynamical response to the
altitude of the anomalous radiative tendencies. All in all, our results
suggest that solar-induced changes in ozone, both in the
tropics/mid-latitudes and the polar regions, are important for modulating
the SH dynamical response to the 11-year solar cycle. In addition, the
markedly nonadditive character of the SH polar vortex response simulated in
austral spring highlights the need for consistent model implementation of
the solar cycle forcing in both the radiative heating and photolysis
schemes.
Introduction
It is now well understood that changes in the incoming ultraviolet (UV)
radiation associated with the 11-year solar cycle influence temperatures and
ozone concentrations across much of the stratosphere (e.g. Penner and
Chang, 1978; Brasseur and Simon, 1981; Haigh, 1994; Randel et al., 2009;
Ramaswamy et al., 2001; Keckhut et al., 2005; Soukharev and Hood, 2006;
Mitchell et al., 2015b; Maycock et al., 2016). In addition to being a major
driver of decadal variability within the stratosphere, these effects can
initiate a dynamical response that propagates down into the troposphere
(e.g. Kuroda and Kodera, 2002; Kodera and Kuroda, 2002), thereby affecting
surface climate variability (e.g. Thíeblemont et al., 2015). The
incoming UV radiation (increased for enhanced solar cycle activity) is
absorbed in the middle atmosphere by oxygen and ozone molecules, the
photolysis of which (Eqs. 1, 3) leads to formation of ozone,
predominantly in the stratosphere, and shortwave heating (Eq. 2):
1O2+hν(λ<242nm)→O+O,2O+O2+M→O3+M(M=N2,O2,…),3O3+hν(λ<1180nm)→O+O2.
Clearly, the heating of the stratospheric air parcels through direct
absorption of solar radiation by ozone and the photochemical production of
ozone are closely coupled. However, this is not necessarily the case in
atmospheric models. In chemistry–climate models (CCMs), shortwave heating
from ozone is usually handled by the radiation scheme, a crucial physical
component of any climate model. A photochemistry module in turn solves the
chemical reactions that lead to ozone production. The accuracy of individual
schemes, as well as the method for implementing the solar cycle forcing, can
vary substantially between models (e.g. SPARC, 2010; Sukhodolov et al.,
2016). Such differences are likely to affect the simulated responses to the
solar cycle forcing across different CCMs. Furthermore, not all climate
models include an interactive chemistry module and, therefore, are capable
of including a feedback from ozone that is consistent with the imposed
spectral solar irradiance (SSI) changes and the resulting adjustments of
temperature and transport. In general, there has been a wide spread of
modelled atmospheric responses to the 11-year solar cycle forcing reported
in the literature (e.g. Austin et al., 2008; SPARC, 2010; Mitchell et al.,
2015a; Hood et al., 2015). A number of these multi-model studies have
attempted to attribute the spread of modelled atmospheric responses to the
solar cycle forcing to the details of specific aspects of model design
(e.g. the resolution of the radiation scheme and the height of the model top); such a task is, however, inherently difficult owing to the wide
diversity in model design.
In the spirit of understanding the contributions of modelled radiation and
photolysis processes to the simulated 11-year solar cycle response, this
paper examines the responses to the amplitude of the 11-year solar cycle
with the forcing included separately in either the radiation or photolysis
scheme. While some studies reported results of similar calculations made
with fixed dynamical heating (FDH) models (e.g. Shibata and Kodera, 2005;
Gray et al., 2009) or for only the annual mean using a CCM (e.g. Swartz et
al., 2012), separating the impacts of these processes on the 11-year solar
cycle response at seasonal timescales has not, to our knowledge, received
much attention in the literature. Clearly such decomposition is, by
definition, an idealised study owing to the strong physical coupling between
the radiative and photochemical processes in the atmosphere. However, this
is a valuable exercise as it helps to elucidate the factors that can affect
the modelled response to the 11-year solar cycle forcing and thus whether
these may contribute to the divergent multi-model results described above.
We focus here on the direct responses to the solar cycle forcing in the
tropics (yearly mean), as well as on the corresponding circulation responses
in the Southern Hemisphere (SH) during winter and spring. It is now well
established that the SH high-latitude stratosphere experiences on average
lower wave activity than the Northern Hemisphere (NH). This makes the SH
polar vortex stronger, less variable on interannual timescales and closer to
the thermodynamical equilibrium than its NH counterpart, thereby enhancing
the detection of the solar-induced anomalies in the region. In addition,
while the solar-induced dynamical response in the NH, including its
underpinning mechanisms, has received considerable attention in the
literature (e.g. Yukimoto and Kodera, 2007; Ineson et al., 2011; Scaife et
al., 2013; Andrews et al., 2015; Gray et al., 2016), the corresponding SH
dynamical response and the mechanisms driving it are not as extensively
examined (e.g. Haigh and Roscoe, 2006; Kuroda and Shibata, 2006; Kuroda et
al., 2007; Petrick et al., 2012; Kuroda and Deushi, 2016).
Section 2 discusses the model and experiments used. Section 3 introduces the
yearly mean temperature responses to the amplitude of the 11-year solar
cycle with the forcing included exclusively in either the radiation or
photolysis scheme, as compared to the control case that includes them both,
and points out the key regions discussed in this paper. Section 4 discusses
the tropical yearly mean responses in the simulations performed, and Sect. 5 discusses the corresponding SH dynamical responses in winter and spring. This is
followed by a consideration of a potential explanatory mechanism for the
different effects of the solar cycle forcing in the photolysis and radiation
schemes (Sect. 6) and the discussion of the results (Sect. 7). The paper
is summarised in Sect. 8.
The experiments
We use the United Kingdom Chemistry and Aerosol Model coupled to version 7.3
of the Met Office Unified Model (UM-UKCA) in the atmosphere-only HadGEM3-A
r2.0 configuration (Hewitt et al., 2011). The chemistry scheme used is the
extended chemistry of the stratosphere scheme (CheS+), as described in
Bednarz et al. (2016). Unlike in Bednarz et al. (2016), however, the model
version used here does not include the coupling of stratospheric aerosols
with the radiation and photolysis schemes.
The implementation of the 11-year solar cycle forcing in the radiation and
photolysis schemes is identical to that described in Bednarz et al. (2019).
The yearly mean total solar irradiance (TSI) data used are those recommended
for the CMIP5 (Coupled Model Intercomparison Project 5) simulations
(Fröhlich and Lean, 1998; Lean, 2000, 2009; Wang et al., 2005),
processed to force the mean of the 1700–2004 period to be 1365 W m-2
(Jones et al., 2011). A fit to spectral data from Lean et al. (1995) is used by the
radiation scheme to account for the change in partitioning of solar
radiation into wavelength bins. In the Fast-JX photolysis scheme used here
(Telford et al., 2013), the change in partitioning of solar irradiance into
wavelength bins is accounted for by scaling the photolysis bins according to
the difference in the yearly mean CMIP5 SSI data for the years 1981 and 1986
(solar maximum, SMAX; and solar minimum, SMIN) and the
long-term evolution of the processed TSI. A more detailed description of the
implementation of the 11-year solar cycle variability in UM-UKCA, including
an evaluation of the atmospheric response to the 11-year solar cycle
forcing, can be found in Bednarz et al. (2019). Note that since the model
uses prescribed sea surface temperatures (SSTs), the full tropospheric response to the imposed change
in solar forcing will not be captured, as tropospheric temperatures are
strongly constrained by the imposed SSTs.
Long time series are needed in order to confidently diagnose the atmospheric
response to the 11-year solar cycle forcing. Therefore, in order to increase
the signal-to-noise ratio while minimising the computational requirements of
long transient integrations, a number of perpetual-year time-slice
integrations are performed under either perpetual SMAX or SMIN conditions.
These are represented by the annual mean solar forcing conditions for the
years 1981 and 1986, respectively (ΔTSI = 1.06 W m-2). All
other forcings are climatological and identical in all runs. These include
the 1977–1987 mean of the SSTs and sea ice (Rayner et al., 2003), as well as of surface and aircraft emissions of CO, HCHO (both surface-only) and NOx following the CCMVal2 (Chemistry-Climate Model Validation 2) specifications
(Morgenstern et al., 2010). The levels of greenhouse gases and ozone-depleting substances for the year 1982 are used according to the SRES A1B scenario (IPCC, 2000) and WMO (2011), respectively. Ozone (as well as N2O, CH4, CCl3F, CCl2F2, C2Cl3F3 and CHClF2) in all runs is treated interactively; i.e. the chemical
ozone field, transported by the circulation, is also used by the radiation
scheme.
We present the results from six 50-year-long (+10 years spin-up)
integrations combined into three SMAX/SMIN pairs. The first pair,
INTERO3SMAX/SMIN, represents the control case with the 11-year solar
cycle forcing implemented consistently in both the radiative heating and
photolysis schemes. In the second pair, RAD-ONLYSMAX/SMIN, the solar
cycle forcing is implemented exclusively in the radiative heating scheme. In
the photolysis scheme, no solar cycle modulation of the spectral
distribution is used in either SMAX and SMIN, but note that the indirect
impact on ozone through changes in atmospheric temperatures and transport
will be captured. The third pair, PHOT-ONLYSMAX/SMIN, is analogous to
RAD-ONLYSMAX/SMIN, but the solar cycle forcing is included exclusively
in the photolysis scheme while constant TSI and SSI are used in the
radiation scheme. Importantly, as noted above, the perturbed ozone field
from the photochemistry is passed to the radiation scheme and will therefore
couple back onto climate. We analyse the resulting differences between the
simulated SMAX and SMIN responses for each pair, and, for brevity, henceforth
we refer to them without any subscripts as INTERO3, RAD-ONLY and PHOT-ONLY.
The experimental set-up is summarised in Table 1.
Summary of the sensitivity time-slice experiments performed.
Yearly mean zonal mean temperature change (K) between SMAX and
SMIN for (a) RAD-ONLY, (b) PHOT-ONLY and (c) INTERO3. Hatching in (a)–(c)
shows regions where the response is statistically significant at the 95 %
level (calculated using a two-tailed Student t test). Shown also (d) is
the difference (shading) between the sum of the single-forcing responses and
INTERO3 (contours, as in c).
The yearly mean temperature response
Figure 1 shows the simulated yearly mean SMAX–SMIN temperature responses in
the single-forcing experiment pairs (RAD-ONLY and PHOT-ONLY, a–b) and in the
control pair with both forcings included (INTERO3, c). In RAD-ONLY, the
temperature response maximises near the tropical and mid-latitude
stratopause at ∼0.4–0.5 K. In PHOT-ONLY, the response
simulated in this region is somewhat smaller (up to ∼0.3–0.4 K); its magnitude also decreases less rapidly with decreasing altitude. With
the exception of a small overestimation in the tropical lower mesosphere,
the response obtained by combining the single-forcing responses in the
tropics agrees with the response in the control pair (up to ∼0.6–0.7 K, d).
Importantly, the individual responses to direct radiative heating and
photolysis cannot be linearly combined to capture the total response in the
high latitudes, in particular in the SH. The stratospheric temperature
increase in INTERO3 decreases slowly at latitudes poleward of 60∘
in both hemispheres (Fig. 1c). In contrast, PHOT-ONLY shows a distinct
yearly mean warming of the SH polar stratosphere (up to ∼0.6 K). The magnitude of this polar temperature response exceeds that found near
the tropical stratopause. In comparison, the yearly mean temperature in
RAD-ONLY does not change substantially throughout most of the Antarctic
stratosphere. The sum of the yearly mean RAD-ONLY and PHOT-ONLY responses
(RAD+PHOT) over the Antarctic shows up to ∼0.5 K
difference compared to the INTERO3 response (Fig. 1d). This is large enough to exceed
the ±2 standard error confidence interval around the INTERO3 response
(not shown), although we note that the difference between RAD+PHOT and
INTERO3 responses is not significant in a strict statistical sense when the
confidence intervals around both RAD+PHOT
Where the standard
errors in PHOT-ONLY and RAD-ONLY are added in quadrature.
and INTERO3 are
considered.
We therefore concentrate in this paper on two regions: firstly the tropics,
where the stratospheric responses appear mostly linearly additive; and
secondly the SH high latitudes, where they do not.
Yearly mean tropical average (25∘ N–25∘ S)
change in (a) the shortwave heating rates (K d-1), (b) temperature
(K), and (c) ozone (%) between SMAX and SMIN for RAD-ONLY (red),
PHOT-ONLY (blue) and INTERO3 (black), together with the associated
confidence intervals (±2 standard errors). The green line indicates
the sum of the RAD-ONLY and PHOT-ONLY responses.
The tropical yearly mean responseShortwave heating rates
Figure 2a shows the yearly mean tropical mean (25∘ S–25∘ N) SMAX–SMIN differences in shortwave heating rates (SWHRs) in the three
pairs of experiments. In RAD-ONLY, the SWHRs increase directly due to the
increased solar radiation and the resulting enhanced absorption by ozone. In
PHOT-ONLY, even though the prescribed SSI does not change in the radiative
scheme calculations, the increased levels of ozone (Sect. 4.3; Fig. 2c)
enhance the SWHR, as described by Haigh (1994).
The maximum amplitudes of the tropical mean SWHR responses in the two
single-forcing pairs of experiments, ∼0.08–0.09 K d-1
near the stratopause, are not distinguishable from one another based on the
estimated uncertainties, and, thus, both effects contribute almost equally to
the maximum SWHR anomaly near the stratopause. The RAD-ONLY tropical
response is largest at ∼50–60 km and then decreases sharply
with decreasing altitude within the stratosphere. This is related to the
intensity of UV radiation being attenuated with increasing path length
through the atmosphere. In comparison, the PHOT-ONLY response is smaller
than in RAD-ONLY above ∼60 km (not shown; see also top of
Fig. 2a) but significantly larger in the mid-stratosphere (e.g. by a factor
of two at ∼40 km). This is due to the SMAX–SMIN increase in
tropical ozone in PHOT-ONLY that maximises in the mid-stratosphere
(∼37 km, Fig. 2c). Thus, while the contributions from the
photolysis and radiation schemes to the SWHR changes are similar near the
stratopause, the impact of the enhanced photochemical production of ozone
dominates in the mid-stratosphere (in agreement with Shibata and Kodera,
2005, and SPARC, 2010).
The tropical mean SWHR response in INTERO3 reaches up to ∼0.16 K d-1 and mostly follows the sum of PHOT-ONLY and RAD-ONLY
(green line in Fig. 2a). Thus, in the tropics, the individual SWHR
responses in the single-forcing experiments can be added linearly to give an
estimate very close to the full response.
Temperature
The corresponding SMAX–SMIN tropical average temperature responses are shown
in Fig. 2b (where ΔTSI = 1.06 W m-2). In INTERO3, the maximum temperature response peaks at ∼0.6 K over a fairly broad
layer spanning ∼45–60 km. It is noteworthy that, despite the identical
implementation of the 11-year solar cycle forcing in the model, the maximum
response simulated in these time-slice runs is somewhat smaller than the
response found in the analogous transient UM-UKCA integrations discussed in
Bednarz et al. (2019, ∼0.8 K W-1 m2), likely indicating
some contributions of indirect dynamical processes and/or interannual
variability to one or both responses. In both cases, the UM-UKCA-simulated
temperature response is somewhat smaller than that found in some reanalyses (e.g.
Mitchell et al., 2015b; Bednarz et al., 2019); this could be due to large
uncertainties in the responses diagnosed from reanalyses and/or some
deficiencies in the model implementation of the solar cycle forcing (see
Bednarz et al., 2019, for details).
Our integrations show significant SMAX–SMIN changes in the upper
stratospheric temperatures in RAD-ONLY and PHOT-ONLY, illustrating that the
solar cycle impacts on both atmospheric heating and photolysis are important
in determining the temperature response there. As noted earlier, there is a
large spread in the simulated upper stratospheric temperature responses to
the 11-year solar cycle forcing among different atmospheric models (e.g.
Austin et al., 2008; SPARC, 2010; Mitchell et al., 2015a; Hood et al., 2015).
Thus, details of both schemes in models and their implementation of the
solar cycle forcing can have a strong influence on the simulated
stratospheric temperature response to the 11-year solar cycle and thus
contribute to the inter-model spread.
The estimated standard errors in the magnitude of the temperature responses
are comparatively larger than those found for the SWHRs, presumably owing to
the additional contribution from dynamical processes to the stratospheric
temperature variability through adiabatic heating and cooling. Thus, the
temperature responses in RAD-ONLY and PHOT-ONLY are statistically
indistinguishable throughout most of the stratosphere. We note that although
PHOT-ONLY shows a somewhat stronger SWHR response in the upper stratosphere
than RAD-ONLY (Fig. 2a), the associated PHOT-ONLY temperature response there
is smaller (Fig. 2b). This illustrates that the atmospheric temperature
response to the amplitude of the 11-year solar cycle forcing is not only
controlled by changes in SWHRs, but also reflects the associated changes in
the longwave component as well as any indirect changes in the circulation
(not shown). As discussed above, the combined RAD+PHOT stratospheric
temperature response in the tropics is in good agreement with the results
from INTERO3 (consistent with Shibata and Kodera, 2005; Gray et al., 2009;
and Swartz et al., 2012).
Ozone
Figure 2c shows the simulated changes in the tropical mean ozone mixing
ratios. In RAD-ONLY, we find a small SMAX–SMIN ozone decrease (up to
∼0.5 %) in the mid- to upper stratosphere and lower
mesosphere. This results from the enhancement of chemical ozone loss under
increased temperature, most importantly through the Chapman and NOx
ozone loss cycles (Fig. S1 in the Supplement; with the change in ozone loss via
the Chapman cycle being a factor of ∼1.5–6 larger between
40 and 50 km than via the NOx cycle; see also e.g. Barnett et al., 1975;
Haigh and Pyle, 1982; Jonsson et al., 2004). In contrast, ozone increases in
PHOT-ONLY throughout most of the stratosphere and lower mesosphere. This
occurs primarily due to the enhanced photolysis of oxygen at wavelengths
shorter than ∼242 nm (Eq. 1) and the subsequent formation of
ozone (Eq. 2), but the response is also influenced by a solar-induced reduction in the
stratospheric NOx levels (not shown), likely related to its enhanced
photochemical removal (e.g. Sukhodolov et al., 2016). The maximum tropical
mean stratospheric ozone response in PHOT-ONLY (∼3 %) is
somewhat larger than in INTERO3 (∼2.5 %.), reflecting the
inverse dependence of ozone on the associated temperature changes (with the
temperature-induced modulation of the NOx cycle playing the dominant
role in the mid-stratosphere, Fig. S1; see also
Jonsson et al., 2004). Throughout most of the tropics, the yearly mean
RAD+PHOT ozone response is in a reasonable agreement with the response in
INTERO3 (in agreement with Swartz et al., 2012). There is some
overestimation of the summed response compared with the control case; this
illustrates that stratospheric ozone concentrations are controlled by a
range of photochemical processes, thereby resulting in a complex dependence
of the SMAX–SMIN ozone response on the associated temperatures, incoming
wavelength-dependent solar radiation and any resulting changes in
ozone columns above.
To summarise, in the tropics the SMAX–SMIN changes in the SWHRs, temperature, and ozone in PHOT-ONLY and RAD-ONLY, which include the solar cycle forcing
only in the photolysis and radiation schemes, respectively, can be summed
linearly to give a response that is in a good agreement with the full
response in the control INTERO3 pair. Our UM-UKCA results agree with the
previous FDH calculations of Shibata and Kodera (2005), Gray et al. (2009)
and SPARC (2010), as well as with the CCM results of Swartz et al. (2012).
However, as noted in Sect. 3, the results show larger differences between
the combined and the control temperature responses at high southern
latitudes (Fig. 1d). The following section analyses the corresponding
responses modelled during the SH winter and spring, where the role of
dynamical processes in modulating the response to solar cycle forcing has
been shown to be important (Kuroda and Kodera, 2002; Kodera and Kuroda, 2002).
The seasonal response in the Southern Hemisphere
The mechanism proposed by Kuroda and Kodera (2002) and Kodera and
Kuroda (2002) (thereafter referred to as KK2002a and KK2002b) to explain the
dynamical response to the 11-year solar cycle forcing they identified in
reanalysis data postulates that solar-induced changes in the tropical SWHRs
and temperatures initiate a chain of feedbacks that modulates the strength
of the polar vortex during the dynamically active season. The UM-UKCA-simulated changes in zonal mean zonal wind and temperature during SH winter
(June–August, JJA) and spring (September–November, SON) for the three pairs
of experiments are shown in Figs. 3 and 4, respectively.
The SMAX–SMIN differences in zonal mean zonal wind modelled in the SH high
latitudes in INTERO3 are fairly weak and not highly statistically
significant in either winter or spring (panels e–f in Figs. 3–4). There is a suggestion of a weak (∼0.5 m s-1) strengthening of
the polar vortex near the stratopause during winter, consistent with the
strengthened horizontal temperature gradient. In comparison, the reanalysis
data suggest a strengthening of the SH polar jet on its equatorward side and
weakening on its poleward side in winter; this spatial pattern is followed
by an enhanced weakening/warming of the vortex in austral spring (e.g.
KK2002a; KK2002b; Frame and Gray, 2010; Mitchell et al., 2015b; Kodera et
al., 2016). The disagreement between the model results and reanalysis data
could be due to a number of factors, including (i) the uncertainties in the
reanalysis SH response; (ii) differences between the time-slice runs here with
prescribed climatological SSTs/sea ice and a transient evolution of the real
atmosphere and its coupling to the oceans; and (iii) a positive bias in the model
SH zonal wind climatology (not shown), which may affect interactions between
planetary waves and the mean flow.
In RAD-ONLY, the zonal mean SH zonal winds in winter strengthen between SMAX
and SMIN on the equatorward flank of the stratospheric/lower mesospheric jet
by up to ∼3 m s-1 (Fig. 3a). This is associated with a
cooling of the high-latitude stratosphere (up to ∼0.75 K, Fig. 4a). The strengthening of the polar vortex in the mid-latitudes
extends down to the extratropical troposphere, where it is accompanied by a
small (∼0.5 m s-1) negative zonal wind anomaly in the
subtropical troposphere. The latter is indicative of a small poleward shift
in the mid-latitude eddy-driven jet (Haigh et al., 2005; Simpson et al.,
2009). Whilst the modelled stratospheric responses in RAD-ONLY are generally
not highly statistically significant, they bear some resemblance to those
found in reanalysis studies (e.g. KK2002a; KK2002b; Frame and Gray, 2010;
Hood et al., 2015; Mitchell et al., 2015b; Kodera et al., 2016). No
significant high-latitude response was simulated in RAD-ONLY in austral
spring (panel b in Figs. 3–4).
In contrast, in PHOT-ONLY there is a strengthening of the stratospheric jet
on its poleward side (up to ∼1 m s-1) and a weakening on
its equatorward side (up to ∼2.5 m s-1) during SH winter
(Figs. 3c and 4c). This represents a poleward contraction of the polar
vortex and is accompanied by a warming in the mid- to upper high-latitude
stratosphere of up to ∼1 K. Importantly, the easterly zonal
wind anomaly develops with time, with significantly weaker zonal wind (up to
∼3.5 m s-1) simulated in the SH mid- to high-latitude
upper stratosphere and lower mesosphere in spring (Fig. 3d). Coincident with
the zonal wind changes, the Antarctic stratosphere is warmer by up to
∼2 K in the austral spring (SON) mean (Fig. 4d). This
modulation of the polar vortex persists until the vortex breaks up. A
histogram showing the interannual variability of the mid-latitude zonal
winds in August simulated in all runs is shown in Fig. S2.
Shading: seasonal mean (left: JJA, and right: SON) SH zonal mean
zonal wind change (m s-1) between SMAX and SMIN for (a, b) RAD-ONLY, (c, d) PHOT-ONLY and (e, f) INTERO3. Single and double hatching indicate statistical significance at the 90 % and 95 % confidence level, respectively (t test). Contours show the corresponding climatological seasonal mean zonal mean zonal wind for the respective SMIN run; contour spacing is 10 m s-1.
As in Fig. 3 but for the SMAX–SMIN zonal mean temperature changes
(K) (shading) and climatological zonal mean temperature in SMIN run
(contours). Contours spacing is 20 K (shown for values beginning at 140 K).
The poleward shift of the stratospheric vortex simulated during winter in
PHOT-ONLY and its equatorward strengthening in RAD-ONLY are essentially
opposite to one another. Therefore, there is a substantial cancellation
between the responses upon linear addition of the JJA means. The combined
RAD+PHOT temperature and zonal wind responses in JJA are generally similar
to the weak response in INTERO3 (Fig. 5).
(a) Shading shows the JJA mean difference between the sum of the
RAD-ONLY and PHOT-ONLY temperature (K) responses and INTERO3. (b) As in
(a) but for the corresponding difference in zonal wind (ms-1) responses. Contours in
(a, b) show the responses in INTERO3 for reference.
As in Fig. 5 but for the SON mean.
In austral spring (SON), a different picture emerges for the comparison
between the sum of the PHOT-ONLY and RAD-ONLY responses and the INTERO3
response. In particular, the development of a significantly weaker and
warmer polar vortex in PHOT-ONLY in spring contrasts strongly with the small
circulation changes found in RAD-ONLY and INTERO3. Consequently, the
combined RAD+PHOT responses in SON show larger differences compared to
INTERO3 (Fig. 6). There is a difference in polar temperatures of up to
∼1.75 K between the summed and INTERO3 responses, which is
large enough to exceed the ±2 standard error confidence interval
around the INTERO3 response and be evident in the annual mean (Fig. 1d). We
note, however, that this SON difference between RAD+PHOT and INTERO3 is
not significant in a strict statistical sense where the confidence intervals
around both responses are considered. Nonetheless, our UM-UKCA results
highlight that stratospheric high-latitude dynamical responses to the
amplitude of the 11-year solar cycle forcing are complex and could be
nonadditive. We explore this behaviour next.
Proposed mechanism for the non-linear SH springtime response
The mechanism for the 11-year solar cycle modulation of the polar vortex
proposed by KK2002a/b centres on the direct solar-induced warming in the
tropical region in autumn/early winter and the immediate changes in the
horizontal temperature gradients as the primary driver of the chain of
feedbacks between planetary waves and the mean circulation throughout the
winter season. To understand the potential reasons for the different dynamical
responses simulated among our UM-UKCA experiments, we focus here on the
changes in the SWHRs, the primary driver of the anomalous temperature
tendencies. We use a simple measure for the solar-induced changes in the
horizontal SWHR gradient across the SH as given by Eq. (4):
ΔSMAX-SMINSWHRgradient=ΔSMAX-SMINSWHR0-60∘S4-ΔSMAX-SMINSWHR60∘S-90∘S.
SH spring
First, we look at the reasons behind the non-linear springtime response. The
original mechanism proposed by KK2002a/b considers only the direct
solar-induced temperature changes in the tropics during autumn/early winter
as the primary driver of the high-latitude circulation responses throughout
the dynamically active season. However, our results suggest that changes in
the SWHR gradients throughout the whole time period are important for the
evolution of the SH dynamical response. During spring it is the SWHR changes
at higher latitudes, influenced strongly by the changes in ozone, that can
be particularly important for determining the horizontal gradients owing to
the increasingly higher mean insolation following the onset of spring.
In particular, the springtime changes in the SWHR horizontal gradients near
60∘ S in RAD-ONLY and INTERO3 (Fig. 7b) have similar vertical
structure and both are much smaller than their corresponding gradient
changes in winter (Fig. 7a). These small SON horizontal gradient changes,
arising from the similarity between SWHR responses in the
tropics/mid-latitudes and the polar regions (Fig. S3b, d), give rise to small zonal wind and temperature responses in the
two pairs of experiments. In stark contrast to this, PHOT-ONLY shows a
markedly different SWHR gradient change (Fig. 7b): while the gradient
strengthens substantially at ∼40 km, the response is negative
in both the lower mesosphere and in the lower stratosphere. The gradient
response is dominated by the strong contribution of the high-latitude SWHR
response changes, which show an alternating positive and negative pattern (Fig. S3d).
These high-latitude SWHR changes are strongly related to the changes in
polar ozone (Fig. 8). We find that ozone mixing ratios in PHOT-ONLY increase
in winter and spring not only in the tropics but also throughout large parts
of the polar stratosphere (Fig. 8c–d). In fact the percentage changes in
polar ozone, in particular during spring, can be larger than those in the
tropical/mid-latitude stratosphere. These are likely to occur due to a
combination of elevated ozone levels already locally present before the
start of the dynamically active season (not shown) and changes in the
circulation/transport. In line with the simulated enhancement of the
stratospheric meridional circulation (Fig. 9) and, thus, increased
transport of ozone-rich air from the tropics and higher polar altitudes,
ozone anomalies are transported poleward and downward; the percentage ozone
anomalies also appear to magnify in spring. Further feedbacks may also be
possible due to any resulting coupling with temperature and/or chemical loss
cycles that may follow (e.g. Hood et al., 2015). As more solar radiation
reaches the high latitudes in late winter/spring, any changes in ozone there
become increasingly important for determining the horizontal SWHR gradients
and, hence, for feeding back and modulating the mean flow. This marked
pattern of changes in SWHR gradients in PHOT-ONLY accompanies comparatively
larger zonal wind and temperature responses (Figs. 3 and 4). The schematic
representation of such a mechanism is in Fig. 10b and d.
SH winter
Secondly, we consider why the two single-forcing experiment pairs (RAD-ONLY
and PHOT-ONLY) indicate contrasting SH winter polar vortex responses. We
find that during winter (JJA) the maximum changes in the SWHR gradient near
60∘ S (Fig. 7a) in our runs peak at different altitudes, with the
strongest changes in gradient found in the lower mesosphere in RAD-ONLY and
in the upper stratosphere in PHOT-ONLY. Little insolation reaches the SH
high latitudes in winter, and thus the SWHR responses there are small (Fig. S3c), so that the changes in the horizontal
gradients in winter are dominated by the SWHR responses in the SH
tropics/mid-latitudes (Fig. S3a). The latter are largely similar to those
found for the tropical annual mean in Fig. 2a, following the same arguments
as in Sect. 4.1. We also find that the development of the SH zonal wind and
temperature anomalies in our experiment pairs is associated with changes in
planetary wave propagation and breaking: the wave propagation/breaking is
increased in PHOT-ONLY and reduced in RAD-ONLY, with no well-defined changes
in INTERO3 (Figs. S5 and S6). To our knowledge few
studies have examined the role of the spatial structure of the anomalous
solar-induced tropical temperature tendencies for the resulting high-latitude dynamical response (e.g. Ito et al., 2009, who looked at horizontal
structure). Possibly, the propagation and breaking of planetary waves within
the stratosphere may be sensitive to the spatial, in our case the vertical,
structure of the anomalous SWHRs. These would act to alter temperature
tendencies, thereby influencing zonal winds and potential vorticity
gradients that are important for planetary wave propagation. The details of
such potential sensitivity are, however, difficult to diagnose using our
experiments and this hypothesis requires further examination with additional
sensitivity experiments. Another potential reason for the differences in the
simulated winter responses between the integrations may be the role of
zonally asymmetric ozone heating in influencing planetary wave propagation.
Numerous studies have shown that stratospheric ozone, as a radiative gas,
can influence planetary wave propagation, thereby impacting on the
interaction between the planetary waves and mean flow (e.g. Nathan and
Cordero, 2007; Kuroda et al., 2007, 2008; McCormack et al., 2011;
Silverman et al., 2018). Possibly, the presence of increased ozone levels in
PHOT-ONLY may act in a similar manner, enhancing the impact of such zonally
asymmetric ozone heating. As before, this hypothesis should be subject to
further testing. The schematic representation of the proposed mechanism is
shown in Fig. 10a and c.
The sum of the changes in the SWHR gradients in RAD-ONLY and PHOT-ONLY
agrees with INTERO3 in austral winter. However, this agreement is not found
in austral spring: the sum of the single-forcing responses is dominated by
the changes in PHOT-ONLY and is not additive, consistent with the lack of
additivity of the zonal wind and temperature responses in SON (Fig. 6).
Therefore, our results highlight the need to implement the solar cycle
forcing interactively in both the radiative heating and photolysis schemes
to fully capture the complex feedbacks between the photochemistry, radiation
and dynamics.
Seasonal mean JJA (a) and SON (b) change in the SWHR gradient (K d-1), as defined in Eq. (), between SMAX and SMIN for INTERO3
(black), PHOT-ONLY (blue), RAD-ONLY (red), and for the sum of PHOT-ONLY and RAD-ONLY
(green).
Seasonal mean (left: JJA, right: SON) SH zonal mean changes in
ozone mixing ratios (%) between SMAX and SMIN for (a, b) RAD-ONLY, (c, d) PHOT-ONLY and (e, f) INTERO3. Single and double hatching indicate statistical significance at the 90 % and 95 % confidence level, respectively (t test). Note the additional contours at ±0.5 %.
The monthly mean evolution of the polar cap average
(90–60∘ S) change in the vertical component of the
transformed Eulerian mean circulation, w‾∗ (mm s-1), between SMAX and SMIN for (a) RAD-ONLY, (b) PHOT-ONLY and (c) INTERO3. Positive values indicate anomalous upwelling and vice versa. Thick
white and grey lines indicate statistical significance at the 90 % and
95 % confidence level, respectively.
Schematic representation of the proposed mechanism. Yellow ovals
represent changes to the SWHRs; red and blue ovals represent strengthening
and weakening of zonal mean zonal wind, respectively. The green arrow
indicates changes in ozone along the meridional circulation, the wavy black
arrows the propagation of planetary waves (increased/decreased as given by
the plus/minus signs), and the dotted green line an interaction between
ozone and planetary waves. The dashed horizontal lines indicate tropopause
and stratopause, and the grey areas in (a, c) indicate the regions covered in polar night.
Discussion
Haigh (2010) pointed out that the solar-cycle-induced ozone response alters
the penetration of solar radiation to lower altitudes and, therefore, leads
to a stratospheric SWHRs response that is complex and non-linear, depending
on the associated changes in ozone. In agreement, our results indicate that
the changes in ozone associated with photochemical production and coupling
to the circulation, not only in the tropics/mid-latitudes but also in the
polar regions, are important for modulating the SH dynamical response to the
amplitude of the 11-year solar cycle. A similar conclusion was reached by
Hood et al. (2015), who suggested that both increases in tropical ozone and dynamically induced sharp horizontal ozone gradients at higher
latitudes are important for horizontal temperature gradients in the
stratosphere and thus play a role in amplifying the associated seasonal
zonal wind response. In addition, Kuroda and Kodera (2005), Kuroda and
Shibata (2006) and Kuroda et al. (2008) found that under higher solar
activity any changes in polar ozone driven by the Brewer–Dobson circulation
during winter can persist in the lower stratosphere for several months,
thereby inducing local temperature and zonal wind responses. In agreement,
while the dynamical response simulated in RAD-ONLY largely disappears by
spring, the response in PHOT-ONLY develops with time, with changes in polar
ozone potentially contributing to this behaviour.
The importance of springtime high-latitude ozone changes in modulating the
SH polar vortex has also been recognised in the context of halogen-induced
Antarctic ozone depletion (e.g. McLandress et al., 2010; Keeble et al.,
2014). There have also been indications that the role of ozone, as a
radiatively active gas, is important in influencing the interactions between
planetary waves and the mean flow, thus modulating the dynamical response to
the solar cycle forcing (e.g. Kuroda et al., 2007, 2008; Nathan and
Cordero, 2007). McCormack et al. (2011) showed that the inclusion of
zonally asymmetric ozone heating in their model weakens the climatological
winter NH polar vortex. The idea that increased ozone levels at SMAX may act
in a similar manner has been proposed by other studies (e.g. Kuroda et al.,
2007, 2008; Nathan and Cordero, 2007), although the importance of this
effect for the solar SH dynamical response has been recently questioned
(Kuroda and Deushi, 2016).
Hood et al. (2015) argued that it is important that models reproduce the
significant ozone and temperature responses that have been observed in the
tropical upper stratosphere in order to simulate stronger amplification of
the horizontal temperature gradients at these altitudes. A comparison with
the altitude differences between the changes in the SWHR gradients found in
RAD-ONLY and PHOT-ONLY in winter raises an interesting question of whether
the SH dynamical response could be sensitive not only to the magnitude of
the changes in the SWHR gradient but also to its maximum altitude range. It
is now accepted that variability in the tropical stratosphere can affect the
high latitudes due to its impact on the planetary wave propagation and
breaking. For instance, a number of studies reported evidence for the
influence of the Quasi-Biennial Oscillation (QBO) on the polar vortex or on
the development of the NH high-latitude dynamical response to the solar
cycle forcing (e.g. Holton and Tan, 1980; Labitzke et al., 2006; Ito et al.,
2009; Matthes et al., 2013; Watson and Gray, 2015). Assuming that changes in
the zonal momentum forcing associated with the different phases of the QBO
modulate the vertical structure of the tropical temperatures, then a similar
mechanism involving changes in wave–mean-flow interactions may operate
here, although specially designed experiments would be required to further
diagnose the details of these sensitivities.
All in all, the apparent nonadditive character of the dynamical response
simulated in our experiments during the SH spring argues strongly for the
need to include the solar cycle forcing interactively in both the radiation
and photolysis schemes in order to fully capture the atmospheric response to
the 11-year solar cycle.
Conclusions
The atmospheric response to the amplitude of the 11-year solar cycle forcing
in the UM-UKCA chemistry–climate model has been separated into the
contributions resulting from direct radiative heating and from changes in
photolysis. Pairs of sensitivity time-slice experiments representing maximum
and minimum conditions of the 11-year solar cycle were performed with the
solar cycle forcing included exclusively in either the model radiation or
photolysis scheme. The sum of the two single-forcing responses was compared
with a control pair with both effects included.
In the tropical upper stratosphere, the yearly mean SMAX–SMIN shortwave
heating rate responses in the radiation-only and photolysis-only experiments
were found to be of similar magnitudes, with both resulting in significant
temperature responses near the stratopause. Details of the implementation of
the solar cycle forcing in the individual schemes in models will have an
important influence on the simulated tropical stratospheric temperature
responses to the solar cycle forcing. Hence, this will be important when
considering the large inter-model spread in the atmospheric response to the
11-year solar cycle forcing reported in the literature. Below the
stratopause, the shortwave heating anomaly in the radiation-only case
decreases sharply with decreasing altitude and is smaller than in the
photolysis-only experiment. However, the corresponding upper stratospheric
temperature response is ∼0.1 K larger, illustrating that the
stratospheric temperature response to the amplitude of the solar cycle
forcing is not just the result of the shortwave heating rate perturbation
but is also influenced by any changes in the longwave component as well as
any indirect dynamical processes (Sect. 4.2). For ozone, the
radiation-only case shows a small (∼0.5 %) decrease in the
tropical upper stratospheric ozone at SMAX due to the acceleration of
chemical ozone loss at higher temperatures. In contrast, in the
photolysis-only case tropical ozone abundances increase by up to
∼3 % due to the enhanced O2 photolysis and the
subsequent ozone production. The magnitude of the tropical stratospheric
ozone response in the photolysis-only case is slightly larger than in the
control case, in line with the inverse dependence of ozone concentrations on
temperature.
The pairs of experiments showed different SH high-latitude circulation
responses between the 11-year solar cycle maximum and minimum in austral
winter and spring. In the radiation-only case, the stratospheric responses
at high southern latitudes are not highly statistically significant, but
the results suggest a strengthening of the polar vortex during winter on its
equatorial side and a cooling of the polar stratosphere at solar maximum
broadly consistent with the reanalysis. In contrast, in the photolysis-only
case we find a poleward contraction of the polar vortex and an associated warming
of the polar stratosphere. In JJA, the sum of these two distinct responses
shows strong cancellation and compares well with the small vortex response
in the case including both radiation and photolysis effects together.
However, this agreement was not found in austral spring (SON), where the
springtime weakening and warming of the polar vortex found in the
photolysis-only case is in stark contrast to the negligible responses in
the other simulations.
In order to understand a mechanism behind the different dynamical behaviour
in our runs and the resulting non-linear springtime response, an analysis of
the corresponding shortwave heating rate gradients across the Southern
Hemisphere was performed. We find differences in the magnitude and vertical
structure of their changes in winter. This raises a question about a
potential sensitivity of the dynamical response to the altitude of the
anomalous radiative tendencies, although this hypothesis requires further
testing. Another potential factor contributing to the different winter
responses may be the role of enhanced zonally asymmetric ozone heating
brought about by the increased ozone levels in modulating planetary wave
propagation and breaking, Our results thus act as a motivation for further
study. Importantly, we find marked changes in the Antarctic shortwave
heating rates in the photolysis-only case in spring; these make a strong
contribution to the associated changes in the horizontal shortwave heating
rate gradients. These high-latitude changes are predominantly driven by the
photochemical ozone changes and their coupling to the circulation changes
(Figs. 8 and 9), but further feedbacks due to any resulting coupling with
temperatures/chemical loss cycles could also play a role. As changes in the
horizontal shortwave heating rate gradients throughout the dynamically
active season could feed back on and modulate the mean flow, this is a
plausible mechanism to explain the simulated weakening and warming of the
polar vortex in spring.
All in all, the tropical yearly mean shortwave heating rates, temperature
and ozone responses in both the photolysis-only and the radiation-only cases
are found to be important for determining the full direct stratospheric
response to the amplitude of the 11-year solar cycle forcing, with both
effects being largely, albeit not fully, additive in the tropics. However,
the apparent nonadditive character of the high-latitude dynamical responses
simulated in the SH spring strongly argues for the need to include the solar
cycle forcing interactively in both the radiation and photolysis schemes in
order to capture the complex feedbacks between photochemistry, radiation and dynamics and, thus, in order to fully model the atmospheric response to the
11-year solar cycle forcing.
Data availability
The model output is available on request.
The supplement related to this article is available online at: https://doi.org/10.5194/acp-19-9833-2019-supplement.
Author contributions
EMB ran the model experiments, analysed the data and wrote the paper, with
discussion, feedback and input from ACM, PB and JAP. NLA provided the model
version and PJT implemented the 11-year solar cycle forcing into the model.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We
acknowledge the use of HECToR, the UK's national high-performance computing
service. The authors thank Remi Thiéblemont and one anonymous reviewer
for helpful comments that have improved the paper.
Financial support
We acknowledge funding from the ERC for the ACCI project grant number 267760, including PhD studentship for EMB. ACM, JAP, PJT and NLA were
supported by the National Centre for Atmospheric Science, a NERC-funded
research centre. ACM acknowledges support from an AXA Postdoctoral
Fellowship and a NERC Independent Research Fellowship (NE/M018199/1).
Review statement
This paper was edited by Jens-Uwe Grooß and reviewed by Rémi Thiéblemont and one anonymous referee.
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