Introduction
The Atlantic Meridional Overturning Circulation (AMOC) is an important
component of climate variability in the North Atlantic region
. The surface branch of the AMOC transports
heat from the Southern Hemisphere (SH) and the tropics towards the north, is
closely connected to the Atlantic Multidecadal Oscillation, and contributes
to the temperate climatic conditions in western Europe .
Thus, understanding the variations in the strength of this circulation is
important in particular for future climate change
, decadal climate
predictions , and with respect to potential
abrupt climatic changes as proposed for the past
.
(a) Total solar irradiance (TSI) anomaly of
-3.5 Wm-2 (dashed) and -20 Wm-2 (solid) applied in
this study. (b) Global annual mean ensemble mean 2 m temperature.
(c) Ensemble mean AMOC index in the different experiments, smoothed
using a 5-year running mean. Dots denote significant differences in the
(un-smoothed) annual mean values between the SRR ensemble and the control
ensemble (Student's t test, p ≤ 0.05). Small stars below the CHEM
time series correspond to years with significant differences between the CHEM
and NOCHEM experiment (p ≤ 0.05). The beginning and the end of the
SRR period are indicated by vertical lines in panels (b) and
(c).
Several processes are involved in driving the AMOC, ranging from internal
processes of the climate system such as the thermohaline process
to external forcing
. The purpose of this study is to assess the influence of a
reduction in the solar forcing on the AMOC. In particular, we investigate the
role of chemistry–climate interactions in modulating the response of the
atmospheric circulation to reduced solar radiation and their effect on the
AMOC. To this end, we perform ensemble sensitivity simulations for different
solar radiation reductions (SRRs) with a state-of-the-art coupled
atmosphere–ocean–chemistry–climate model, where atmospheric chemistry is
either enabled or disabled.
So far, the external forcing response of the AMOC has been mainly studied in
climate models without interactive atmospheric chemistry .
Thereby, volcanic eruptions have been found to intensify the AMOC on decadal
timescales , through a reduction in sea surface
temperatures and a shift of the North Atlantic Oscillation (NAO) towards its
positive phase. Moreover, volcanic eruptions may excite the variability in
the AMOC . The response, however, may depend on the
background conditions of the climate system . An
increase in the solar forcing has been found to weaken the AMOC by increasing
sea surface temperatures (SSTs) and enhancing freshwater input
and has been proposed
to be a driver of Greenland temperature variations
.
Among many possible influences, a change in the solar forcing may also affect
the NAO . For example,
the circulation in the polar stratosphere during winter (polar night jet) has
been proposed to be affected by a change in the ultraviolet (UV) radiation
. By stratosphere–troposphere interactions, stratospheric
anomalies can propagate down to the troposphere and cause circulation
anomalies at the surface . A positive phase
of the NAO is then associated with a strengthening of the polar night jet and
vice versa and may also affect
the AMOC .
The stratospheric response to UV variations is modulated by chemistry–climate
interactions . In particular, stratospheric ozone
reacts to solar irradiance changes. The increase in solar UV enhances the
shortwave heating rate through ozone absorption e.g..
Additional solar UV in the Herzberg continuum (λ < 242 nm)
intensifies ozone production, while UV in the Hartley band destroys ozone
e.g.Fig. 1. Because the solar UV variability decreases
with wavelength the first effect prevails and leads to ozone increase in the
middle stratosphere in phase with the increase in the solar UV. In turn the
ozone increase gives additional heating with magnitude comparable to primary
heating by the increase in solar UV alone . This process
can amplify the efficiency of the earlier mentioned top-down propagation
and is obviously missing if the ozone concentration is
prescribed.
Still, most of these studies are based on models without interactive
atmospheric chemistry. The influence of climate changes on the state of the
ozone layer has long been recognized. The cooling of the
stratosphere by greenhouse gases (GHGs) slows down catalytic ozone oxidation
cycles, leading to ozone increase e.g.. The
greenhouse warming accelerates Brewer–Dobson circulation reducing ozone in
the tropical lower stratosphere and enhancing its abundance over middle to
high latitudes . The ozone changes have
substantial implications for the climate. The influence of the ozone recovery
associated with the implementation of the Montreal Protocol limitations on
the production of ozone destroying substances on the SH has been identified
in observations and model simulations
e.g.. Recently, it was suggested that the use
of interactive chemistry instead of prescribed ozone climatology can
influence climate model properties. showed that the
application of interactive chemistry reduces the climate sensitivity by
3–8 %. A similar reduction in the climate sensitivity was also found by
. A more substantial reduction in the model response to
4 × CO2 by up to 20 % due to taking into account interactive chemistry was
reported by . All these studies attributed the reduction to
the changes in ozone, water vapour, and clouds in response to climate warming.
These conclusions were not confirmed by very recent results of the CESM1-WACCM
model , which found a similar ozone response to 4 × CO2 but no
changes in climate sensitivity. In contrast, applied the
same model and demonstrated that the interactive ozone introduces a negative
feedback leading to a weaker surface warming due to an enhancement of the
solar irradiance. Thus, these results show that further experiments are
necessary in order to assess the model discrepancies and to deepen our
understanding of the ozone feedback and its importance for the simulation of
future climate change under the influence of different natural and
anthropogenic factors.
The outline of this study is as follows. The model configuration and the
experiments are described in Sect. 2. Section 3 presents the results, first
for the experiments without interactive atmospheric chemistry and then through an
analysis of the differences causes by the chemistry–climate interactions. A
summary and concluding discussion is given in Sect. 4.
Model and experiments
The model
We use the coupled atmosphere–ocean–chemistry model SOCOL-MPIOM to simulate
the effect of a change in solar activity on the climate
. SOCOL consists of the atmospheric
component ECHAM5 coupled to the chemistry module MEZON
. The middle-atmospheric configuration of
ECHAM5 is used , which resolves the atmosphere up to
0.01 hPa (about 80 km) with 39 levels. The horizontal
resolution is T31, corresponding to a grid size of
3.75∘ × 3.75∘.
The chemistry is directly coupled to ECHAM5 and uses temperature data to
calculate the tendency of 41 gas species, taking into account 200 gas-phase,
16 heterogeneous, and 35 photolytical reactions. Optionally, the coupling to
MEZON can be disabled. In this case a three-dimensional time-dependent ozone data
set needs to be specified.
The shortwave radiation scheme of SOCOL considers spectral solar irradiance
(SSI) values in six spectral bands. Time series for each spectral interval
are used as forcing to allow for changes in the spectral composition of the
total solar irradiance. The shortwave scheme considers Rayleigh scattering,
scattering on aerosols and clouds, and the absorption of UV by O2,
O3, and 44 other species. Additional parametrizations for the
absorption of UV in the Lyman-alpha, Schumann–Runge, Hartley, and Higgins
bands are implemented following and . The
longwave scheme considers wavenumbers between 10 and
3000 cm-1 and takes into account water vapour, CO2,
O3, N2O, CH4, CFC-11, CFC-12, CFC-22, aerosols, and
clouds.
With the given vertical resolution, SOCOL is not able to produce a
Quasi-Biennial Oscillation (QBO). Thus, a QBO nudging is applied
. The time step of the atmospheric component is 15 min, with the full radiation and chemical computations updates
performed every 2 h.
SOCOL is coupled to the ocean model MPIOM
using the OASIS coupler . MPIOM includes an
embedded sea ice module. To avoid numerical singularities at the North Pole,
both poles of the rotated Arakawa C grid are shifted and placed over land
(Greenland and central Antarctica). The nominal resolution is 3∘ –
varying between 22 and 350 km – with a higher resolution in the deep
water formation regions in the North Atlantic and the Weddell Sea. Convection
is implemented by greatly enhanced vertical diffusion when the water column
becomes unstable. Sea ice dynamics are based on the viscous–plastic rheology
formulated by . A constant sea ice salinity of 5 psu is
assumed. The time step of the oceanic component is 2 h and 24 min.
The experiments
Ensemble sensitivity simulations with SOCOL-MPIOM are performed to study the
effect of SRR on the climate system and the AMOC.
Such SRRs are caused by either a grand solar minimum or solar radiation
management techniques. Ten simulations are carried out for each ensemble
experiment; the experiments differ in the solar forcing applied and whether
or not chemistry–climate interactions are considered in the model.
Perpetual AD 1600 conditions and zero volcanic aerosols (i.e. excluding the
volcanic eruption of Huaynaputina) are applied in all simulations. For the
sensitivity simulations only the solar forcing is allowed to change in time.
The solar forcing consists of the SSI and photolysis rates.
As reference experiment we perform two control ensembles, CTRL_CHEM and
CTRL_NOCHEM, with and without interactive chemistry, respectively. In these
experiments all forcings represent the conditions of the year AD 1600,
including the solar forcings of the year AD 1600. The year 1600 was chosen
since a stable long-term control simulation with SOCOL-MPIOM was available
from previous studies . Note that the
differences in the climatic conditions between 1600 and the commonly used
year 1850 are small and both represent a pre-industrial climate state.
The two SRRs simulated for this study are characterized by a step-wise total
solar irradiance (TSI) reduction of -3.5 and -20 Wm-2,
referred to as S1 and S2, respectively (Fig. a). The S1 SRR is
comparable to a grand solar minima like the Dalton minimum or Maunder minimum
in a large-amplitude solar forcing reconstruction
e.g.. With -20 Wm-2 the S2 SRR is
comparable to a weak solar radiation management scenario ,
which may counteract an increase in the radiative forcing from GHGs of about
3 Wm-2. The reduction in the solar forcings is switched on at year
5 of a simulation and lasts for 30 years when it is switched off and the
simulation is continued for 25 years. Both SRRs are simulated with and
without interactive chemistry and are named S1_CHEM, S2_CHEM, S1_NOCHEM,
and S2_NOCHEM in the following. A summary of the experiments performed for
this study is given in Table .
Overview of the ensemble experiments used in this
study. Each ensemble consists of 10 experiments.
Experiment
TSI (Wm-2)
Chemistry
CTRL_CHEM
const.
Yes
CTRL_NOCHEM
const.
No
S1_CHEM
-3.5
Yes
S1_NOCHEM
-3.5
No
S2_CHEM
-20
Yes
S2_NOCHEM
-20
No
In the CHEM experiments, ECHAM5 and MEZON are coupled and the atmospheric
chemistry responds to the solar radiation changes. In NOCHEM, temporal and
spatial ozone variations need to be prescribed. Therefore, a daily 3-D ozone
climatology is applied, based on a AD 1600 control simulations.
All ensemble simulations are initialized from model year 1300 of a long
control simulation with interactive chemistry performed under perpetual
AD 1600 conditions . The ensemble members only differ in
their initial conditions by slightly perturbing the atmosphere (atmospheric
restarts for 1, 2, 3 January, etc.). The oceanic component is always
initialized using the same initial conditions.
Note that we erroneously applied a slightly different solar forcing in 6 of
10 simulations. This TSI difference of 0.018 Wm-2 is caused by
a different rounding of the SSI values and leads to very small differences
between the control ensemble experiments and the SRR experiments already
prior to the start of the reduction.
The AMOC index is calculated by selecting the maximum in the annual mean
meridional overturning streamfunction northward of 28∘ N and below
300 m. To detect the influence of the stratospheric circulation on
the troposphere and the AMOC we use the hemispheric mode of the Northern
Hemisphere (NH) the Arctic Oscillation (AO). While the NAO is more closely
related to the AMOC, the AO has a stronger imprint of stratosphere
troposphere interactions. The AO index is defined as the spatially averaged
monthly mean sea level pressure difference between 40 and
65∘ N, which is normalized by the mean and the standard deviation of
the corresponding control ensemble. Furthermore, the index is multiplied by
-1 to reflect the negative phase of the AO by negative values and vice
versa. Using a different definition of the AO (based on empirical orthogonal function or using the sea
level pressure north of 70∘ N) or an index of the NAO leads to very
similar results.
Results
Both SRRs lead to a significant reduction in the global mean near-surface
(2 m) air temperature (Fig. b). For the stronger S2 experiment
the cooling is more pronounced than for the S1 experiments and reaches -1.0
and -0.9 K for S2_CHEM and S2_NOCHEM, respectively (averaged over
the last 5 years of the SRR period). For the S1 experiment, the temperatures
reduce by -0.1 K in both ensembles. The temperature
instantaneously responds to the imposed radiation drop and reaches the lowest
values at the end of the reduction period. The continuous cooling in the
course of the SRR, which is well visible in the S2 ensembles, suggests that
the model has not yet reached thermal equilibrium. In fact, from the model's
equilibrium climate sensitivity (for a doubling of CO2, see
) an equilibrium temperature response of -1.3 K
is expected for S2_CHEM and -1.4 K for S2_NOCHEM. However,
a comparison with the CO2 sensitivity is only a rough estimate, since the
climate sensitivity (and the contributions from chemistry–climate
interactions) differs between the solar and CO2 forcing and depends on the
sign of the forcing perturbation .
The larger cooling in the CHEM experiments is related to differences in the
stratospheric response. In particular, stratospheric ozone concentrations are
reduced due to the reduced UV radiation (Fig. S1a, d in the Supplement),
a process which is not considered in the NOCHEM experiments. Additionally,
water vapour concentrations are affected by the SRR. In S2_NOCHEM, the
largest anomalies (-15 %) are found in the tropical upper
troposphere, but stratospheric reductions exceed -10 % almost
everywhere (Fig. S1c in the Supplement). In S2_CHEM, the stratospheric
reductions in water vapour are more pronounced (up to -35 %), due
to the effect of the solar forcing on the oxidation of methane, the most
important in situ source of stratospheric water vapour (Fig. S1b in the Supplement). Due to the
greenhouse effect of ozone and water vapour, the outgoing longwave flux
increases more in CHEM than in the NOCHEM and leads to an additional cooling
of the troposphere. The positive water vapour anomalies found in the
uppermost model levels in the CHEM experiments (Fig. S1b, e in the
Supplement) are related to the reduced UV photolysis of the water vapour
molecules.
A slight initial reduction in the global mean temperature is also found in
the reference ensemble experiments and is related to the initial conditions
of the ocean. With all ensemble simulations sharing the same oceanic
conditions in the beginning, the AMOC development of the first years is
dominated by the oceanic memory. During the first decade of the experiments
a decline in the AMOC from 21.0 to 19.8 Sv is found
(Fig. c). This decline is very similar in both reference
experiments. The minimum state of the AMOC is reached in year 12–13 of the
reference experiments and in the following years the AMOC increases to its
maximum value of 21.4 Sv in year 35.
Annual mean 2 m temperature anomalies
(colours), sea level pressure anomalies (red contours), and 50 % sea
ice extent line (yellow contours) averaged over the SRR period. Temperature
and sea level pressure anomalies are calculated relative to the control
ensemble mean; for the sea ice extent, the values of the control ensemble and
the S2 experiments are depicted by the solid and dashed line, respectively.
Panel (a) shows the difference for the S2_CHEM ensemble, and the S2_NOCHEM
anomalies are shown in (b). Panel (c) displays the
differences between S2_CHEM and S2_NOCHEM. Dots denote non-significant
temperature differences (Student's t test, p > 0.05). The sea level
pressure contour interval is 0.25 hPa and negative sea level pressure
anomalies are dashed.
(a–i) S2_CHEM (a, d, g),
S2_NOCHEM (b, e, h), and the difference between S2_CHEM and
S2_NOCHEM (c, f, i) ensemble mean upper ocean (0–220 m) density
anomalies (kg m-3) for late winter (January–March) averaged over the
first (a–c), second (d–f), and last
decade (g–i) of the SRR period. Cyan contours display
the extent of the 50 % sea ice area for the CTRL ensemble mean (solid line)
and the SRR experiments (dashed line). Panels (j–r): S2_CHEM
(j, m, p), S2_NOCHEM (k, n, q), and the difference between
S2_CHEM and S2_NOCHEM (l, i, r) ensemble mean January–March mixed
layer depth anomalies (m, shading) averaged over the first
(j–l), second (m–o), and last
decade (p–r) of the SRR period. Contours denote the
average January–March mixed layer depth in CTRL_CHEM and
CTRL_NOCHEM, with a contour step of 500 m. Dots
denote non-significant density or mixed layer depth differences (Student's
t test, p > 0.05).
The AMOC is not affected by the SRR during the first few years of the
simulation. Starting with simulation year 10, however, and even more
pronounced in the second half of the reduction period, the AMOC is
significantly stronger in S1_NOCHEM during several years and in S2_NOCHEM
for most of the years between years 15 and 35 of the experiment. In the CHEM
ensemble simulations no significant AMOC intensification is found for S1. In
S2_CHEM, the AMOC is significantly stronger during the second half of the
SRR period, but the intensification is weaker in comparison to S2_NOCHEM.
The differences between the AMOC index for S2_CHEM and S2_NOCHEM are also
reflected in the anomaly pattern of the AMOC (Fig. S2 in the Supplement). Within the first
10 years the intensification of the circulation is weak. Positive anomalies
are found between 40 and 65∘ N and between the surface and a depth of
2800 m depth. During these first 20 years of the reduction period the
intensification is slightly larger in S2_CHEM. A pronounced strengthening of
the circulation occurs in the second decade of the reduction period. Positive
anomalies cover all latitudes from the Equator to 65∘ N and most
levels between the surface and 3000 m depth. In the second decade the
intensification is more pronounced in S2_NOCHEM. In the third decade,
finally, a further intensification is found, which is again stronger in
S2_NOCHEM. In the following, we will first address the relevant processes
that are responsible for the AMOC intensification (Sect. )
before we assess the role of chemistry–climate interactions in order to explain the
lower sensitivity of the AMOC to SRR in the CHEM experiments
(Sect. ).
The thermal effect of SRR on the AMOC
A direct effect of SRR is the reduction in shortwave energy reaching the
troposphere and the surface and thus in temperature, which is apparent almost
everywhere in the NH (Fig. ). Averaged over the 30-year
reduction period the sea ice growth in the Barents Sea is stronger in
S2_CHEM than in S2_NOCHEM (Fig. ). Furthermore, a larger
cooling over the Barents Sea is found in S2_CHEM, which extends towards
northern Eurasia. In the S1 experiments temperature and sea ice anomaly
patterns are weaker but similar to S2 and S1_CHEM is characterized by an
amplified cooling as well (not shown). During the first 10 years, when no
AMOC differences between the CHEM and NOCHEM experiments are found, the
temperature and sea ice anomalies are very similar. The Arctic sea ice
differences between CHEM and NOCHEM, which emerge in the last 20 years of the
reduction period, are therefore related to the weaker AMOC in the CHEM
experiments and the reduced heat transport into the Arctic.
Temperature and salinity averaged over the upper 220 m
for two deep water formation regions: the North Atlantic and Nordic Seas. The
deep water formation regions cover all grid cells with an annual mean mixed
layer depth ≥ 250 m in the corresponding ocean basins. The lines show
the salinity and temperature development from the beginning (triangle) to the
end (large dot) of the SRR for the S2_CHEM (orange) and S2_NOCHEM (blue)
experiments. Each point represents a single year. To improve visibility, the
values are smoothed using a 15-year low-pass filter. Error bars denote the
mean and the standard deviation of the corresponding control ensembles.
Contours represent the water density.
Annual mean zonal mean temperature
(a–c) and annual mean zonal mean zonal wind
(d–f) anomalies in the S2 experiments relative to the control
experiments: panels (a) and (d) show the anomalies for the
S2_CHEM experiment and (b) and (e) the results for
S2_NOCHEM. The differences between both experiments (S2_CHEM – S2_NOCHEM)
are shown in (c) and (f). Anomalies are averaged over the
30-year SRR period. Contours represent the mean state in the control
experiments with contours from 180 to 280 K (contour step
15 K) for the temperatures and -30 to 30 ms-1 (contour
step 5 ms-1) for the zonal wind. Dots denote non-significant
temperature differences (Student's t test, p > 0.05).
The cooling in the lower atmosphere has a direct effect on the ocean. With a
reduction in the upper ocean temperatures and an increased salinity due to
the enhanced sea ice formation, the density of the upper ocean increases
almost everywhere (Fig. a–i). Additionally, a shift of the
storm track and a significant reduction in the precipitation in the North
Atlantic contributes to the salinity and density increase (not shown). During
the first 10 years of the SRR period, differences in the density anomalies in
the upper ocean of the North Atlantic are small and not significant, except
for a region south of Greenland, where the density is significantly higher in
S2_NOCHEM (Fig. a–c). In the following decade further
increases in the upper ocean density are found in both experiments, but the
anomalies are again larger in S2_NOCHEM (Fig. d–f). During this time, the
density anomalies in large parts of the North Atlantic are more pronounced in
S2_NOCHEM in comparison to S2_CHEM. Finally, in the last 10 years, density
anomalies are still strongly positive, but the differences between both
experiments weaken (Fig. g–i).
Convection takes place in the Nordic Seas and in a region in the North
Atlantic close to the Labrador Sea (contours in Fig. e–h). The
intensity of the deep water formation in these two regions is an important
driver of AMOC variability . Focusing on the changes in
the Nordic Seas, we find an intensification of the deep water formation
already for the first 10 years of the reduction period
(Fig. j–l). A further intensification is found for the second
and the third decade, but the anomalies between S2_CHEM and S2_NOCHEM show
only weak significance. The anomalies in the S1 experiments are similar –
i.e. differences are mostly non-significant. Density changes in the Nordic
Seas are driven by a combination of temperature and salinity changes
(Fig. ). The temperature changes, however, dominate in the first
half of the SRR period, while the increasing salinity drives the density
changes in the second half.
In the North Atlantic the density and mixed layer differences between
S2_CHEM and S2_NOCHEM are larger than the ones in the Nordic Seas. During
the first 10 years of the SRR period, positive mixed layer depth anomalies
are found in S2_NOCHEM (Fig. k), while no consistent response
is found in S2_CHEM (Fig. j). Consequently, the intensification
is significantly stronger in S2_NOCHEM (Fig. l). A similar
picture emerges for the second decade (Fig. m–o). In the third
decade a clear intensification is obvious in S2_CHEM, while a slight
reduction is found in S2_NOCHEM in the southern region of the North Atlantic
convection zone (Fig. p–r). As in the Nordic Seas, the density
changes are driven by the reduced temperatures in the first half of the SRR
(Fig. ). In the second half of the SRR period the salt content
of the upper ocean increases, while temperatures increase again, related to
the intensification of the overturning. The salinity changes, nevertheless,
lead to a further increase in the density in the second half of the
reduction period.
(a, b) Ensemble mean number of
sudden stratospheric warming (SSW) events per winter season (November–March) as in defined by . (c, d) Box plot statistics for the number of SSW events per winter season averaged
over the SRR period. The beginning and the end of the SRR period are indicated
by vertical lines in panels (a) and (b).
The increasing density and deep water formation in both convective regions
help to understand the intensification of the AMOC in the course of the SRR.
Driven directly by the temperature response to the reduced solar forcing,
this mechanism can be considered as the thermal effect of the SRR on the
overturning. However, in S2_CHEM the intensification of the convection in
the North Atlantic is delayed in comparison to S2_NOCHEM. Similar
differences are found between the two S1 experiments (Fig. c).
A further mechanism is therefore needed to understand the differences in the
AMOC response between the CHEM and NOCHEM experiments.
The dynamical effect and the role of chemistry–climate interactions
Chemistry–climate interactions are most pronounced in the stratosphere
e.g.. In particular, the different response of the
stratospheric ozone and water vapour between CHEM and NOCHEM (Fig. S1 in the
Supplement) leads to large differences in the stratospheric temperatures. For
S2_CHEM temperature anomalies of up to -28 K are found in the
upper stratosphere (Fig. a). Above 1 hPa the maximum
cooling is found in the polar latitudes, with a second maximum in the tropics.
In the lower and middle stratosphere, the cooling is stronger in the tropics
and mid-latitudes. With about -10 K, the maximum cooling in
S2_NOCHEM is much smaller than the response in S2_CHEM
(Fig. b, c). Furthermore, as a consequence of the missing
response of the ozone concentrations to the reduced solar forcing, the effect
of the lower- and middle-stratospheric cooling on the meridional temperature
gradient is weaker.
The response of the zonal mean wind in the stratosphere agrees well with the
temperature anomalies. For S2_CHEM, a pronounced weakening of the NH and SH
polar vortices is found (Fig. d–f). Using the zonal mean wind
component at 60∘ N and 10 hPa as an index for the intensity of
the NH polar vortex , a reduction of
-43 % is found in S2_CHEM during the winter season (November–March)
when averaged over the SRR period. The largest wind anomalies occur during
the vortex maximum in January. The reduction in S2_NOCHEM is much weaker
(-8 %) than in S2_CHEM. Furthermore, the duration of the winter
period with predominant westerly wind is reduced in S2_CHEM by
-30 % and in S2_NOCHEM by -5 %, respectively, when
defining the start of the winter period by the day with the first occurrence
of a westerly daily mean zonal mean wind component at 60∘ N and
10 hPa after September and the end by the first day with easterly winds
after March. Qualitatively similar results are found for the S1 experiments,
with November–March vortex anomalies of -9 % for S1_CHEM and
-2 % for S1_NOCHEM (Fig. S3 in the Supplement). These responses
highlight the non-linear relationship between the solar forcing and the
atmospheric dynamics.
The weakening of the NH polar vortex is closely related to the occurrence of
sudden stratospheric warming (SSW) events (Fig. ). SSW events are
stratospheric extreme events in which the westerly flow during winter time
is reversed and a strong warming in the polar stratosphere is observed. SSW
events in the NH are associated with a breakdown of the polar vortex.
Following the SSW definition by , almost a doubling of the
number of SSW events is found in S2_CHEM (1.34 events per winter in comparison
to 0.68 events per winter in CTRL_CHEM). In S1_CHEM an increase to 0.73 events
is simulated. Similarly to the NH polar vortex, the effect of the SRR on the
SSW events is small in NOCHEM. For S1 the average number of events increases
from 0.68 events per winter in CTRL_NOCHEM to 0.70 events per winter in S1_NOCHEM.
In S2_NOCHEM an increase to 0.73 events is simulated. While the increase in
the mean number of SSW events is small in S2_NOCHEM, a clear reduction in
the years with a low number of SSW events is found (lower quartile of the
box plot).
(a, b) Ensemble mean AO index
(standardized and reversed sea level pressure difference between
45 and 65∘ N) per winter (November–March). Dots
indicate winters with significant differences to the CTRL ensemble (Student
t test p ≤ 0.05). (c, d) Box plot statistics for
the AO index averaged over the SRR. The beginning and the end of the SRR
period are indicated by vertical lines in panels (a) and
(b).
The NH polar vortex and extreme events like SSW affect the tropospheric
circulation in the NH by stratosphere–troposphere interactions. A downward
propagation of wind speed anomalies from the middle stratosphere to the
surface is related to positive and negative phases of the AO
. For a negative phase of the AO, negative wind anomalies
in the stratosphere occur up to 40 days before the AO event takes place at
the surface (Fig. S4 in the Supplement). For a positive phase of the AO, the zonal wind
anomalies are even stronger (not shown). Overall, the downward propagation of
wind speed anomalies does not differ substantially between the CHEM and
NOCHEM control experiments.
The stratospheric changes in the course of the SRR therefore affect the
tropospheric pressure systems. In Fig. a the sea level pressure
anomalies for S2_CHEM reveal a pattern of positive anomalies over large
parts of the Arctic and negative anomalies in the North Atlantic and the
North Pacific, similar to a negative phase of the AO. In S2_NOCHEM,
comparable negative and positive pressure patterns are found, but the
anomalies are much weaker (Fig. b). Due to the strength of the
response, the winter phenomenon AO is reflected in the annual mean values
(Fig. ). However, when focusing on the winter season (November–March) and the AO index, the strength of the anomalies in S2_CHEM is even
more apparent (Fig. ). During the entire SRR phase a persistent
negative phase of the AO is found in S2_CHEM. In S1_CHEM the tendency
towards a negative AO is found as well, although the response is weaker and
several years with a positive phase of the AO occur during the SRR. In the
NOCHEM experiments the response is in general weaker, but a shift towards a
negative AO phase from CTRL_NOCHEM to S1_NOCHEM and S2_NOCHEM is apparent
(Fig. d). In particular, negative AO phases tend to occur more
often in the first half of the SRR period, while neutral conditions dominate
in the second half.
Atmospheric chemistry–climate interactions therefore lead to pronounced
differences in the dynamical response to the SRR, from the stratosphere down
to the surface of the NH high latitudes. With a shift in the pressure pattern
which affects the wind systems in the lower atmosphere, these differences have
the potential to also modify the oceanic circulation.
Influence of a negative AO phase on different oceanic
variables in CTRL_CHEM during winter (November–March). Linear regression
coefficients for (a) net downward heat flux, (b) sea ice
area fraction, (c) liquid freshwater flux (evaporation minus
precipitation), (d) sea surface salinity, (e) sea surface
temperature, and (f) mixed layer depth. To highlight the influence
of a negative AO phase, the AO index has been reversed in the regression
analysis. Dots denote non-significant temperature differences (Student's
t test, p > 0.05).
The control experiments are used to assess the influence of the AO phase on
the North Atlantic. Regressing the AO index on different oceanic variables
reveals that a negative AO phase is associated with an increased downward
heat flux south of Greenland and negative heat flux anomalies close to the
east coast of North America during winter in CTRL_CHEM
(Fig. a). Sea ice cover is reduced in the Labrador Sea
(Fig. b) and the dynamical changes lead to an increased total
freshwater flux into large parts of the North Atlantic, as well as a reduced flux in
the Nordic Seas (Fig. c). These changes cause a reduction in the
salinity (Fig. d), except for a small region south of Greenland,
which may be affected by a weakening of the East Greenland Current.
Additionally, SSTs increase south of Greenland (Fig. e), related
to the enhanced downward heat flux. Since the density of the water decreases
with increasing temperature and decreasing salinity, all these changes lead
to a pronounced reduction in the mixed layer depth (Fig. f). In
CTRL_NOCHEM the effect of the AO is very similar (Fig. S5 in the Supplement).
These changes at the ocean surface are also reflected in the AMOC index. In
both control experiments the AMOC reacts within the same winter season to the
AO phase, as detected by the positive correlation between the winter AO and
the AMOC index of the same season (Fig. S6 in the Supplement). Furthermore, the AO phase has
long-lasting effect on the overturning, reflected in significant positive
correlations for lags up to 9 years.
To summarize, the weaker intensification of the AMOC in the CHEM experiments,
in comparison to NOCHEM, is related to a second (dynamical) response to the
SRR. With interactive chemistry, the stratospheric cooling is strongly
amplified by stratospheric ozone loss. As a consequence, the weakening of the
northern polar vortex is more pronounced, which has larger effects on the
tropospheric circulation patterns, in particular the phase of the AO. The
dynamical changes decrease the density of the surface ocean waters south of
Greenland, reduce convection, and weaken the AMOC. In the NOCHEM experiments
a tendency towards a negative phase of the AO is found as well, although less
pronounced, due to the absence of chemistry–climate interactions. The
dynamical effect on the AMOC is therefore much weaker and the thermal
response dominates.
Conclusions and discussions
Sensitivity experiments for different solar minima and model configurations
with and without chemistry–climate interactions have been carried out to
study the response of the AMOC to reduced solar forcing and the modulating
role of chemistry–climate interactions. Without interactive chemistry the
response of the AMOC is dominated by the direct thermal effect, leading to an
intensification of the overturning circulation. A second dynamical effect is
identified in the experiments with chemistry–climate interactions and leads
to a weakening of the overturning.
Flowchart summarizing the thermal and dynamical
effect of a change in solar radiation on the AMOC. The signs indicate the
correlation between two processes. Dashed boxes represent effects which are
amplified by chemistry–climate interactions.
The two processes are summarized in Fig. : the thermal effect
is related to the reduced shortwave energy reaching the troposphere and the
surface and the ensuing cooling of the lower atmosphere and the upper ocean.
This increases the sea surface density and enhances convection. The thermal
effect, however, is compensated for by the dynamical effect when atmospheric
chemistry is taken into account. Induced by the reduction in the tropical
stratospheric temperatures, a weakening of the NH polar vortex and – by
interactions between the stratospheric and tropospheric circulation –
a negative phase of the AO is found in response to the SRR. The circulation
changes in the troposphere in turn cause a weakening of the AMOC by anomalous
heat and freshwater fluxes. The dynamical effect is amplified by chemistry–climate interactions, due to the enhanced stratospheric temperature response
related to the effect of the reduced UV radiation on the ozone
concentrations. For the weaker S1 SRR, both effects cancel each other out and
therefore no AMOC intensification is found in the experiments with
interactive chemistry. In the S2 experiments with stronger forcing, however,
the thermal response of the AMOC dominates and the dynamical effect leads
only to a reduced intensification of the overturning.
The thermal effect of solar radiation changes on the overturning has been
identified in earlier studies
. Related to
increasing global greenhouse gas concentrations and associated surface
warming, it is also one of the dominant mechanisms for the projected future
weakening of the AMOC
. The
thermal response to the reduced solar forcing also has implications for the
projected weakening of the AMOC in the 21st century. Several studies suggest
that the Sun may enter a grand solar minimum within the next 100 years
, although the amplitude of the
TSI changes is associated with large uncertainties. While the influence on
the global mean temperature increase is small
, the thermal effect may reduce the
projected 21st century AMOC weakening. This is confirmed by experiments of
. The AMOC is significantly stronger in the late 21st
century in ensemble simulations including a grand solar minimum in comparison to ensemble simulations
without a decline in the solar activity (Fig. S7 in the Supplement).
Parts of the dynamical effect have been reported in previous studies. The
relationship between solar variability and the stratospheric circulation was
found for the 11-year cycle as well as for
grand solar minima . The projection of the
stratospheric anomalies on the AO has also been reported in previous studies
. Finally, the influence of the AO
phase on the overturning has been studied
and a few studies have
identified a possible influence of the stratospheric circulation on the
overturning .
The stratosphere responds very rapidly to the reduced solar forcing and the
tropospheric AO index shifts to a negative phase in the second winter after
the onset of the reduction period, although it takes about 5 years before a
persistent negative AO phase is found in S2_CHEM. The response of the AMOC,
however, is delayed by several years. A similar delay was reported by
, who performed sensitivity experiments with an ocean model
forced by different atmospheric conditions. In one experiment a persistent
positive phase of the NAO is simulated and the AMOC responds to this forcing
with strengthening of the circulation, which is delayed by 5–7 years
see Fig. 3 in. This lag of the response agrees
with our results, although an exact timing is difficult to estimate from our
setup.
The influence of the dynamic effect on the AMOC may furthermore depend on the
length of the solar reduction period. found a gradual
weakening of the subpolar gyre response with time in ocean model simulations
forced with a persistent negative phase of the NAO. Additionally, the
response of the AMOC may be non-linear and an increase in the solar forcing
may change the dynamic effect .
Recently, assessed the role of the interactive chemistry
on the temperature and precipitation response to increasing SSI. They
identified a reduced sensitivity with interactive chemistry due to the effect
of the ozone increase on the shortwave radiation balance. Our results for a
SSI reduction indicate a slightly larger temperature sensitivity with
interactive chemistry owing to the effect of the stratospheric water vapour
and ozone changes on the longwave radiation balance. These differences may
be attributed to model differences or differences in the response of the
climate system to increasing and decreasing solar forcing. A possible effect
of the differences in the atmospheric response on the AMOC is not discussed
by .
Here, we show for the first time how stratospheric processes modulate the
modelled response of the AMOC to solar forcing and identify the importance of
chemistry–climate interactions for the response. Hence, previous studies
without atmospheric chemistry may overestimate the sensitivity of the AMOC to
solar forcing since the dynamical effect is absent.
Furthermore, our results reveal possible additional side effects of the solar
radiation management technique: a reduction in the incoming solar radiation
in space to mitigate the temperature increase caused by the emission of GHGs
might affect the tropospheric circulation patterns in the NH and cause a
weakening of AMOC with climatic consequences, in particular for the temperate
climate in western Europe. The dynamical effect is expected to change,
however, when the solar radiation is reduced in the Earth's atmosphere, for
instance, by stratospheric sulfate aerosols. In this case, a strengthening
of the NH polar vortex and a positive phase of the AO may develop, analogous
to the response to strong tropical volcanic eruptions
. This
effect of the positive AO phase may, in turn, lead to an intensification of
the AMOC. Future studies should address the influence of stratospheric
sulfate geoengineering on the AMOC and the possible role of
chemistry–climate interactions.