The stratospheric ozone layer plays a key role in atmospheric thermal
structure and circulation. Although stratospheric ozone distribution is
sensitive to changes in trace gases concentrations and climate, the
modifications of stratospheric ozone are not usually considered in climate
studies at geological timescales. Here, we evaluate the potential role of
stratospheric ozone chemistry in the case of the Eocene hot conditions.
Using a chemistry–climate model, we show that the structure of the ozone
layer is significantly different under these conditions (4×CO2 climate
and high concentrations of tropospheric N2O and CH4). The total
column ozone (TCO) remains more or less unchanged in the tropics whereas it
is found to be enhanced at mid- and high latitudes. These ozone changes are
related to the stratospheric cooling and an acceleration of stratospheric
Brewer–Dobson circulation simulated under Eocene climate. As a consequence,
the meridional distribution of the TCO appears to be modified, showing
particularly pronounced midlatitude maxima and a steeper negative poleward
gradient from these maxima. These anomalies are consistent with changes in
the seasonal evolution of the polar vortex during winter, especially in the
Northern Hemisphere, found to be mainly driven by seasonal changes in
planetary wave activity and stratospheric wave-drag. Compared to a
preindustrial atmospheric composition, the changes in local ozone
concentration reach up to 40 % for zonal annual mean and affect
temperature by a few kelvins in the middle stratosphere.
As inter-model differences in simulating deep-past temperatures are
quite high, the consideration of atmospheric chemistry, which is
computationally demanding in Earth system models, may seem superfluous.
However, our results suggest that using stratospheric ozone calculated by
the model (and hence more physically consistent with Eocene conditions)
instead of the commonly specified preindustrial ozone distribution could
change the simulated global surface air temperature by as much as 14 %. This
error is of the same order as the effect of non-CO2 boundary conditions
(topography, bathymetry, solar constant and vegetation). Moreover, the
results highlight the sensitivity of stratospheric ozone to hot climate
conditions. Since the climate sensitivity to stratospheric ozone feedback
largely differs between models, it must be better constrained not only for
deep-past conditions but also for future climates.
Introduction
The absorption of incoming solar ultraviolet (UV) radiation by stratospheric ozone is responsible for the heating up of the stratosphere and
hence its dynamical stability. In addition, this absorption is essential to
the development of life because it prevents this very harmful UV radiation
from reaching the Earth's surface. Stratospheric ozone is thus a key
component of the radiative equilibrium and habitability of the Earth
(Brasseur and Solomon, 2005). However, although deep-time climates are more
and more investigated with numerical climate models, the role of the
stratosphere in such climates is usually neglected (e.g., Kageyama et al., 2017; Lunt et al., 2017).
The present-day stratosphere has been intensively studied to understand,
anticipate and mitigate the global ozone depletion caused by the emissions
of anthropogenic halogenated compounds such as chlorofluorocarbons and
halons in the second part of the 20th century (WMO, 2014). The phasing out
of the emissions has led to the start of a stratospheric ozone recovery
since the end of the 1990s (Chipperfield et al., 2017). However, in the
context of increasing levels of greenhouse gases (GHGs, e.g., CO2, N2O and CH4) and associated climate change, the
sensitivity of stratospheric ozone to other drivers, especially
climate-related drivers, is increasingly investigated. For example,
stratospheric ozone is sensitive to changes in N2O, CH4 and water
vapor levels. N2O enters in the stratosphere at the tropical
tropopause and controls the levels of NOx, which are the most efficient
ozone-destroying radicals in the middle stratosphere. Enhanced CH4
levels increase ozone production in the troposphere and lower stratosphere
but also lead to higher water vapor levels, which tend to favor ozone
destruction (Bekki et al., 2011; Revell et al., 2012). An increase in CO2
concentration results in a cooling of the stratosphere, which slows down
ozone destruction in the upper stratosphere and hence favors ozone recovery
in this region. In addition to stratospheric chemical changes, the ongoing
climate change tends to intensify the large-scale stratospheric overturning
circulation (the so-called Brewer–Dobson circulation), which is responsible
for upward transport of air in the tropics and poleward and downward
transport at middle and high latitudes (see, e.g., Butchart, 2014, and
references therein). These circulation changes result in reduced ozone
levels in the tropical lower stratosphere due to the faster ascent of air in
the lower tropical stratosphere (Avallone and Prather, 1996) and enhanced
ozone levels at middle and high latitudes (Bekki et al., 2011). This
illustrates how stratospheric ozone responds to climate change. More
recently, Chiodo et al. (2018) presented an analysis of the stratospheric
ozone layer response to an abrupt quadrupling of CO2 concentration in four chemistry climate models. As found previously (see, e.g., WMO, 2014), they showed that increased CO2 levels in the four models lead to a decrease in ozone concentration in the tropical lower stratosphere and an increase at
high latitudes and in the upper stratosphere. However, there were large
differences between models in the magnitude of the ozone response.
At the same time, stratospheric ozone changes also influence the climate.
For instance, climate models have to account for the formation of the
stratospheric ozone hole to be able to reproduce correctly the trends in
Antarctic surface temperatures observed during the last half century (e.g.,
Son et al., 2010; McLandress et al., 2011). Considering larger climate
perturbations, Nowack et al. (2015) performed an abrupt 4×CO2
experiment with a comprehensive ocean–atmosphere–chemistry–climate model and
found that neglecting stratospheric ozone changes triggered by CO2
increase (i.e., specifying a fixed ozone climatology in the model) led to
the overestimation of the surface global mean temperature response by about 1 K (i.e.,
20 % of the total surface temperature response). Chiodo and Polvani (2017) carried out the same numerical experiment (4×CO2) with a
different chemistry–climate model and found that, in contrast to the Nowack et
al. (2015) results, stratospheric ozone changes played a negligible role in the
global surface temperature response. Nonetheless, they found that the
stratospheric ozone feedback in their model significantly reduced the
CO2-induced poleward shift of the midlatitude tropospheric jet by
lowering the strength of the meridional temperature gradient near the
tropopause. These results suggest that stratospheric ozone perturbations
should be accounted for in climate models in order to fully capture the
climate response to GHG changes.
In the past of the Earth, the oxygenated atmosphere has encountered hot
climatic conditions due to strong greenhouse effects. During the early
Eocene (∼56–50 Ma) terrestrial temperatures at high latitudes
were possibly up to 20 K higher than modern ones (Masson-Delmotte et al., 2013). Under such a warm climate, biogenic emissions of N2O and
CH4 were likely to be drastically boosted, being 4 to 5 times higher
than the preindustrial ones (Beerling et al., 2011). Note, however, that, in
most modeling studies of deep-time climates, the role of non-CO2 gases
is neglected (e.g., DeepMIP; Lunt et al., 2017). Beerling et al. (2011)
studied the tropospheric chemical composition under a warm climate and
potentially high biogenic emissions of the early Eocene (55 Ma). Using an
Earth system model including tropospheric chemistry, they found that the OH
concentration, which is the main oxidant for most compounds in the
troposphere, was significantly reduced (by 14 % to 50 %) due to higher
levels of compounds to oxidize. The high tropospheric levels of reactive
greenhouse gases (N2O, CH4 and O3) were maintained under
these conditions. Considering the full Earth system interactions, and in
particular albedo change due to the melting of continental snow, Beerling et al. (2011) calculated an increase of 1.4 to 2.7 K in surface temperatures due to
tropospheric chemical composition changes for the Eocene. However, since
their model did not include stratospheric chemistry, they could not study
stratospheric composition changes. Unger and Yue (2014) investigated the
chemistry–climate feedbacks in the mid-Pliocene (∼3 Ma) using
a vegetation–chemistry–climate model simulating both stratospheric and
tropospheric chemistry. This epoch is cooler than the Eocene but still of
interest because its global climate is thought to be as warm as the climate
projected for the end of the ongoing century (+2–3 K compared to
the present day). Compared to preindustrial conditions (PI), the Unger and Yue (2014) model simulations indicated that the mid-Pliocene ozone burden was
higher by 25 % in the troposphere and by 5 % in the stratosphere. The
global stratospheric ozone increase, resulting from a stronger tropical
upwelling and less ozone destruction in the stratosphere, led to a 20 %
decrease in tropospheric ozone photolysis and hence OH production. As a
consequence, tropospheric OH concentrations were reduced by 20 %–25 % and
hence the lifetime and burden of important reactive species (CO, CH4)
were significantly increased. Unger and Yue (2014) showed that the warming
effect of the changes in chemically reactive compounds (i.e., CH4,
N2O, tropospheric O3) could have represented ∼75 %
of the warming from CO2 increase. The studies of Beerling et al. (2011) and Unger and Yue (2014) suggest that non-CO2 greenhouse gases
may have played a significant role in the overall climate in the Cenozoic
greenhouse worlds.
As pointed out previously, most studies of Cenozoic paleoclimates assume
that the atmospheric composition is fixed except for CO2 because there
is no estimate of these composition changes. The purpose of this paper is
(i) to investigate, using a stratospheric chemistry–climate model, to what
extent the stratosphere, notably the ozone layer, might have been different
in the early Eocene conditions and (ii) to estimate the possible effects of
these stratospheric changes on the tropospheric oxidizing capacity and
climate. The Eocene is characterized by high surface temperatures, elevated
CO2 levels, the absence of ice caps and a large extent of tropical
vegetation. High CH4 and N2O levels are also expected based on Earth system model (ESM) simulations (Beerling et al., 2011). Whereas the data are sparse and have
large uncertainties for the geological past, several proxy-based
reconstructions of CO2 levels and surface temperatures have been
released. They have notably been used to build a harmonized protocol for climate
modeling of the early Eocene (Lunt et al., 2017), which will be used to
intercompare climate model sensitivity during the ongoing Paleoclimate
Modelling Intercomparison Project (PMIP4). This protocol gathers
recommendations on paleogeography, land cover, CO2 and CH4
concentrations, natural aerosols, solar constant, and astronomical parameters, but no recommendations have been provided yet for stratospheric conditions.
In this work, we propose to examine the role of the stratospheric ozone
layer in Eocene climate. We first investigate the stratospheric ozone
response to GHG-induced warm climate such as the one expected under Eocene conditions and compare it to preindustrial climate conditions. We
then discuss the methodology to account properly for these stratospheric
ozone changes in deep-time paleoclimate simulations. Finally, based on the
model simulations, we estimate the difference in UV radiation reaching the Earth surface between this epoch and the preindustrial period and the
resulting impact on tropospheric chemistry. The potential climate forcing of
the stratospheric changes is also discussed.
MethodologyThe LMDz–REPROBUS climate–chemistry model
Simulations are performed with the stratospheric chemistry–climate model
developed in the framework of the IPSL Earth system model (IPSL-CM, climate model)
development (Dufresne et al., 2013). The stratospheric chemistry is computed
with the REPROBUS chemical model (Lefèvre et al., 1994, 1998; Jourdain et al., 2008) coupled with the LMDz atmospheric general circulation model
(Hourdin et al., 2013). REPROBUS describes the chemistry of stratospheric
source gases such as N2O, CH4, CH3Cl and CH3Br and the associated radical chemistry of hydrogen, nitrogen oxides, chlorine and
bromine species. It computes the global distribution of trace gases,
aerosols and clouds within the stratosphere considering gas-phase and
heterogeneous chemistry. The heterogeneous chemistry component takes into
account the reactions on sulfuric acid aerosols and liquid (ternary
solution) and solid (nitric acid trihydrate particles, ice) polar
stratospheric clouds (PSCs). The gravitational sedimentation of PSCs is
simulated as well. The LMDz–REPROBUS chemistry–climate model allows an
interactive coupling of ozone, shortwave heating rates and dynamics as
recommended in Sassi et al. (2005). The resolution of the model is 3.75∘ in longitude and 1.9∘ in latitude, and it has 39 vertical levels, with
around 15 levels above 20 km and around 24 above 10 km and a lid height at
∼70 km.
LMDz and LMDz–REPROBUS have been involved in a large range of studies, model
intercomparisons and evaluations, notably through the participation of the
LMDz model in the international Coupled Model Intercomparison Project (CMIP,
phases 3, 5 and currently 6) and the participation of LMDz–REPROBUS in the Chemistry Climate Model Validation (SPARC CCMVal, 2010) and the Chemistry
Climate Model Initiative (CCMI; Morgenstern et al., 2017). Results presented
in the recent studies of, e.g., de la Cámara et al. (2016a, b),
Thiéblemont et al. (2017) and Ayarzagüena et al. (2018) have shown
that the stratospheric chemistry, dynamics and transport simulated by the
LMDz model and its chemistry–climate model version are consistent with satellite observations,
reanalysis and other models of the same kind.
Simulation setup
The setup of the four simulations performed in this study is summarized in
Table 1. All the simulations consist of 30-year time slices, starting from
atmospheric physical conditions and surface temperatures taken from very
long coupled atmosphere–ocean simulations. For the analysis of our
chemistry–climate simulations, a 5-year spin-up is considered. For all the
simulations, the solar constant is set to 1366 W m-2 and orbital
parameters (obliquity, precession and eccentricity) are set to modern
values as recommended in the DeepMIP protocol (model intercomparison project; Lunt et al., 2017). Oxygen
variations are poorly constrained over pre-Quaternary timescales, and there
is no consensus on the oxygen variations through the Cenozoic (see Fig. 1 of
Wade et al., 2018). In view of these uncertainties, we use a present-day
oxygen content to investigate Cenozoic past climates, as commonly done in
climate models.
Setup of LMDz. AMIP is the Atmospheric Model Intercomparison Project.
Setup nameOzoneCO2CH4N2OSSTPREINDInteractive285 ppm791 ppb275 ppbAMIPEOCENEInteractive1120 ppmExtracted fromEOCENE_OzRoyerPrescribed from RoyerFOAM EoceneEOCENE_Oz1855Prescribed from Szopa et al. (2013)simulation
Note: for the Eocene simulations, the REPROBUS chemical model considers a CH4 concentration of 3614 ppb and an N2O concentration of 323 ppb.
Modeling setup for the Eocene simulations.
Preindustrial simulations
The boundary conditions of our preindustrial experiment (PREIND) include
modern topography, a land–sea mask, ice sheets and climatological mean values
computed over the 1870–1899 period for sea surface temperatures (SSTs) and
sea-ice extent. Greenhouse gases are set to preindustrial values, i.e., a CO2 level at 285 ppm, a CH4 level at 791 ppb and an N2O level at
275 ppb. Halogenated ozone-depleting substances of anthropogenic origin
(i.e., fluorocarbons) are set to zero. Naturally emitted halogenated
compounds (CH3Br and CH3Cl) are prescribed at their preindustrial
levels (respectively, 7 and 482 ppb).
Eocene base case simulation
As for the PREIND experiment, the Eocene experiment (EOCENE) includes
interactive chemistry, which allows us to calculate stratospheric composition.
The physical boundary conditions for the EOCENE experiment (e.g., SSTs,
sea ice, land surface properties) are based on a climate simulation done
with the fully coupled low-resolution Fast Ocean Atmosphere Model (FOAM;
Jacob, 1997) and the Lund–Potsdam–Jena (LPJ) dynamic global vegetation model
(Sitch et al., 2003) coupled offline as illustrated Fig. 1. The LPJ–FOAM coupled
ocean–atmosphere–vegetation simulation provides the surface conditions
(SSTs, land surface conditions) required to simulate the climate with the
LMDz atmospheric general circulation model. FOAM is forced with the Eocene
paleogeography reconstruction of Herold et al. (2014). Compared to the
present-day paleogeography, it includes major modifications, namely closed
Drake and Tasman seaways, an open Central American Seaway, and an open
Paratethys Sea. Topography is altered as well, with a lower Tibetan Plateau
and Andes. CO2 is set to 1120 ppm, equivalent to a 4×CO2 preindustrial level ([CO2]PI), as Eocene CO2 estimates
range between 400 and 2400 ppm (as reported by Lunt et al., 2017, based on
boron isotopes analysis from Anagostou et al., 2016). This CO2 value
lies at the low end of the Eocene-compatible GHG forcing ranges, in
particular those recommended by the DeepMIP project, which proposes to test
3×[CO2]PI, 6×[CO2]PI and 12×[CO2]PI (Lunt et al., 2017). After 2000 model years, SSTs simulated by FOAM are averaged over
the last 100-year period to build a 12-month (seasonally varying)
climatology used as a boundary condition for LMDz. FOAM coupling with the
LPJ vegetation model provides equilibrated vegetation as well, whose
albedo and rugosity are extracted to serve as continental boundary
conditions for LMDz. The global mean SST that we use are 17.3 ∘C for the preindustrial period and 23.9 ∘C for the Eocene. These values lie in
the ranges presented for four model realizations in Lunt et al. (2012). These
ranges are between 15.2 and 17.9 ∘C for PI and between 22.2 and
26.4 ∘C for the Eocene when considering 4×[CO2]PI (Note that
when CO2 varies from 2×[CO2]PI to 16×[CO2]PI, the
range of SST is between 21.4 and 29.7 ∘C; Lunt et al., 2012.) In
addition, the meridional surface temperature gradient is 24.6 ∘C over ocean and 33.7 ∘C over land in our protocol when the ranges with the
4×[CO2]PI experiments shown in Lunt et al. (2012) are [24;
33] ∘C and [25.5; 37] ∘C, respectively. Numerous
recently published paleoclimate studies are based on the two-step
methodology based on FOAM-LPJ and LMDz, and this setup has been shown to
perform well (e.g., Ladant et al., 2014; Licht et al., 2014; Ladant and Donnadieu, 2016; Pohl et al., 2016; Porada et al., 2016; Botsyun et al., 2019).
Applying a coupled vegetation–atmosphere to the particularly warm climate of
the early Eocene (55 Ma), Beerling et al. (2011) have estimated that CH4
and N2O concentrations should have been much higher than nowadays and
could have lain in the 2384–3614 and 323–426 ppb ranges, respectively.
The direct climate impact of highly enhanced CH4 and N2O levels is
accounted for by setting CO2 to a high level (1120 ppm) in the radiative
module in our atmospheric circulation model (Table 1). Ozone chemistry is
affected by changes in N2O and CH4 (e.g., Revell et al., 2012). To
account for this effect, there are CH4 and N2O chemically active
tracers in the REPROBUS chemical model (i.e., modified by the transport and
chemistry schemes). Their surface concentrations are taken from the
modeling study of Beerling et al. (2011), and CH4 and N2O surface
concentrations are set to 3614 and 323 ppb, respectively, in the chemistry
module (REPROBUS). Their global distributions change with time during a
simulation, but they are not used as inputs to the radiative scheme and
hence their changes do not affect the climate; only ozone changes do.
Eocene simulations with prescribed stratospheric ozone
In addition to the EOCENE experiment in which stratospheric ozone is
calculated interactively, two other Eocene simulations
(EOCENE_OzRoyer, EOCENE_Oz1855) are performed
in which different climatological ozone representations are specified
instead of ozone being calculated interactively. The ozone climatology in
the EOCENE_OzRoyer experiment is rather typical of the 1980s
ozone distribution. It originates from fits to the ozone profile from
Krueger and Mintzner (1976) and variations with altitude and latitude of the
maximum ozone concentrations and total column ozone from Keating and Young (1985). This OzRoyer ozone climatology
was constructed by J.-F. Royer (CNRM,
Meteo France) and implemented in the LMDz atmospheric circulation model in
the 1980s. The Oz1855 ozone climatology in the EOCENE_Oz1855
experiment is more representative of the preindustrial period. It is based
on an 11-year mean climatology centered on 1855 derived from historical
transient LMDz–REPROBUS simulations (Szopa et al., 2013). This ozone
climatology is commonly used for the simulation of past climates with the
LMDz model.
The comparison between EOCENE and PREIND experiments, which both include the interactive chemistry of the stratosphere, allows us to explore and
quantify the impacts of the Eocene warm climate on stratospheric circulation
and composition (Sect. 3). Furthermore, comparing the EOCENE experiment
with EOCENE_OzRoyer and EOCENE_Oz1855
experiments allows us to assess the role of the stratospheric ozone
representation on the climate response to Eocene extreme conditions (Sect. 4). Note that the statistical significance of anomaly fields is estimated
here using a Student's t test.
Impacts of Eocene conditions on stratosphereStratospheric ozone in Eocene conditions
We first investigate the impact of Eocene conditions on stratospheric ozone
with respect to preindustrial conditions. Figure 2a and b show the
latitude vs. pressure zonally averaged cross sections of temperature and ozone
anomalies associated with the Eocene conditions. As expected, the CO2
increase leads to a global radiative cooling of the stratosphere with
decreases in temperatures exceeding 12 K above 10 hPa (∼32 km)
and a warming of the troposphere (Fig. 2a). In the troposphere, we further
notice a more pronounced Antarctic amplification of the temperature signals,
which contrasts with present-day climate conditions where a more pronounced
Arctic amplification of the global warming is observed (IPCC, 2013). This
signal, consistently simulated by several models (Lunt et al., 2012), is
linked to the absence of the Antarctic ice sheet in Eocene boundary conditions,
which leads to lower surface topography and albedo. The cooling of the
stratosphere slows down the ozone destruction, resulting in an increase in
stratospheric ozone concentrations (Haigh and Pyle, 1982). This is
consistent with the statistically significant positive ozone anomalies found
above 50 hPa (∼20 km) over the polar regions and above 10 hPa
in the tropical band (Fig. 2b). Note that this effect increases with
altitude in the stratosphere as the photochemical control on ozone levels
becomes prominent (Brasseur and Solomon, 2005). Similarly to the results of
Chiodo et al. (2018), which investigate the effect of quadrupling CO2
by starting from preindustrial climate, a maximum ozone increase of
∼40 % is found at about 2–3 hPa (∼40 km).
Note that in our simulation, the stratospheric chemistry is also modified by
the increase in N2O and CH4. However, their effect only reaches a
maximum of 3 % in the equatorial upper stratosphere (∼5 hPa) (see Fig. S1 in the Supplement). Although this chemical effect on ozone
is statistically significant, its impact appears to be small compared to the
upper stratosphere 40 % increase in ozone due to increasing CO2.
Annual mean differences (EOCENE minus PREIND) of zonally averaged
temperature (in K, a), ozone (in %, b) and age of air (in
years, c). Color-filled contours in (a) and (b) indicate that anomalies
are statistically different at the 1 % confidence level according to a
t test. Black contours show the preindustrial climatology expressed in K (a), ppm (b) and years (c).
In contrast, the lower tropical stratosphere (30∘ S–30∘ N, 100–30 hPa) exhibits a statistically significant ozone decrease of up to
40 %. In this region, the ozone concentration is mostly controlled by
transport processes (Brasseur and Solomon, 2005), especially the strength of
the Brewer–Dobson circulation ascending branch. Figure 2c shows the age of
air (AoA) calculated after 20 years of simulations by taking as a reference
entry point the equatorial lowermost stratosphere, slightly above the
tropopause (i.e., pressure level corresponding to 74 hPa). Globally, the
stratospheric AoA is younger in the Eocene experiment than in the
preindustrial one, revealing an acceleration of the Brewer–Dobson
circulation under Eocene conditions. This, in turn, is consistent with the
reduced ozone concentration in the lower tropical stratosphere. Note also
that the tropopause height is globally lifted up in the Eocene experiment
(not shown). The rise of the tropopause is a robust feature of warmer
climate conditions (Sausen and Santer, 2003) and contributes to the negative
ozone anomaly found in the vicinity of the tropopause region (Fig. 2b)
(Dietmüller et al., 2014).
Next, we examine anomalies of the annual total column ozone (TCO).
Figure 3 shows the comparison of the latitudinal distribution of the annual
TCO for Eocene and preindustrial conditions. In both simulations (Fig. 3a), the TCO shows a minimum in the tropical region (20∘ S,
20∘ N) of ∼270 Dobson units (DU) and maxima near
55∘ N and 55∘ S, followed by poleward decreases that are
more pronounced in the Southern Hemisphere. The differences between Eocene
and preindustrial conditions (Fig. 3b) reveal no changes in the tropical
band 20∘ S–20∘ N but statistically significant positive
anomalies at midlatitudes and in polar regions. The midlatitude maxima
reach ∼390 DU for preindustrial conditions, whereas they
exceed 430 DU for Eocene conditions. This latitudinal distribution of TCO
anomalies is overall consistent with projections of TCO anomalies simulated
in response to the 21st-century climate change and post-CFC (chlorofluorocarbon) era (Li et al., 2009)
or to an abrupt 4×CO2 increase from preindustrial climate conditions
(Chiodo et al., 2018). The detailed comparison of our results with those of
Chiodo et al. (2018) shows, however, noticeable differences at high
latitudes. In our simulations, the TCO anomalies peak at 45∘ S and
50∘ N with maximum differences of 50 and 60 DU, respectively; the
anomalies decrease from these maxima to about 30 DU at high latitudes
(Fig. 3b). In Chiodo et al. (2018), hints of such a decrease were found in
the Southern Hemisphere for only two out of the four models that are
intercompared, and no such decrease was seen in the Northern Hemisphere.
The negative poleward TCO gradient at high latitudes appears to strengthen
markedly under Eocene conditions in the Northern Hemisphere (Fig. 3b). The
seasonal dependence of the TCO high-latitude poleward gradient for the
Eocene and preindustrial conditions in the Northern Hemisphere is explored
in Fig. 4. Figure 4 reveals that, under Eocene conditions, the negative
gradient is particularly pronounced during the winter season (from October
to March), when the stratospheric polar vortex dominates the high-latitude
circulation in the Northern Hemisphere. Hence, this indicates substantial
changes in the stratospheric circulation associated with the Eocene
conditions that we examine further in Sect. 3.2.
Latitudinal profile of the total column ozone (in Dobson units or
DU, a) in the (red) EOCENE and (black) PREIND simulation. Total column
ozone change (in DU, b) between the EOCENE and PREIND simulation.
Dashed lines delimit the 2σ uncertainty envelop, which is
represented by the standard error of the mean.
Zonally averaged seasonal evolution of the latitudinal gradient
(computed as the difference between 84 and 57∘ N) of
the total column ozone for the EOCENE (red) and PREIND (black) simulations.
Dashed lines delimit the 2σ uncertainty envelope, which is
represented by the standard error of the mean.
Seasonal evolution of the Northern Hemisphere stratospheric polar
vortex in Eocene conditions
An overview of the annual average background zonal circulation in
preindustrial conditions and its anomalies associated with Eocene conditions
is shown in Fig. 5. In Eocene conditions, the high-latitude
stratospheric westerlies maxima, indicative of the average location of the
core of the stratospheric Southern and Northern Hemisphere polar night jets
(near 60∘ S and 60∘ N), appear to be overall stronger
and also shifted equatorward in comparison with the preindustrial
climatology (black contour). These results hence suggest a strengthening and
an extension of stratospheric polar vortices, which develop during winter in
each Hemisphere. Note also that the upward extension of the subtropical
upper-tropospheric jets in both hemispheres (centered near 35∘ N/S
around 200 hPa) are consistent with the tropopause rising associated with
Eocene conditions.
Annual mean differences (EOCENE minus PREIND) of zonally averaged
zonal wind (in m s-1). Color-filled contours indicate anomalies that are
statistically different at the 1 % confidence level according to a t test.
Black contours show the preindustrial climatology.
To identify the processes leading to the strengthening of the stratospheric
polar vortex under Eocene conditions, we explore the stratospheric dynamical
wintertime evolution in the Northern Hemisphere, where the largest changes
are found in our simulations (e.g., Fig. 5). Figure 6 shows the monthly
evolution of the zonal-mean zonal wind from October to March in the Northern
Hemisphere. Regardless of simulated climate conditions, the winter season in
the stratosphere is characterized by the development of a mid-to-high-latitude strong westerly jet (or polar night jet – the center of which
roughly corresponding to the edge of the polar vortex), which maximizes in
midwinter (December–January). In early and midwinter (Fig. 6a–e), the
polar night jet in Eocene conditions appears, however, to be twice as strong as
in preindustrial conditions as shown, e.g., in January (Fig. 6d) where the
maximum anomaly near 60∘ N and 5 hPa (∼40 m s-1) associated with Eocene
conditions is larger than the preindustrial climatology (∼30 m s-1). Such
a strong boreal polar night jet was also found in Eocene simulations of
Baatsen et al. (2018). In late winter (Fig. 6f), the differences in polar
night jet strength between the two experiments are no longer statistically
significant in the middle stratosphere and appear to even be reversed in the
upper stratosphere; i.e., an easterly anomaly is found near the stratopause at midlatitudes. This indicates a very fast decay of the polar vortex in Eocene
conditions in late winter (see also Fig. S1 in the Supplement). These
differences in the seasonal evolution of the zonal-mean zonal wind are
consistent with the seasonal evolution of the ozone gradient shown in Fig. 4.
Indeed, under Eocene conditions, the stronger winter polar vortex is
associated with a reinforcement of the mixing barrier at its edge. This leads
in turn to a reduction in air exchanges between mid- and polar latitudes and
hence to a steepening of the poleward ozone gradient. Similar (though less
pronounced) differences between Eocene and preindustrial conditions are found
in the Southern Hemisphere (not shown).
Monthly evolution (October to March) of the zonal-mean zonal wind
differences (m s-1) between the Eocene and preindustrial conditions in the Northern
Hemisphere. Dotted regions indicate that anomalies are insignificantly
different at the 5 % confidence level according to a t test. Black
isolines shows the climatology derived from the preindustrial experiment.
At first glance, the strengthening of the polar vortex under Eocene
conditions may seem inconsistent with the global acceleration of the
Brewer–Dobson circulation as diagnosed by the younger stratospheric age of
air (Fig. 2c). Indeed, a faster Brewer–Dobson circulation is associated
with a stronger planetary wave drag in the stratosphere (i.e., an enhanced
wave breaking), which in turn should lead to a weaker polar vortex. In the
following, we hence investigate the seasonality of the planetary stationary
wave activity and its interaction with the mean flow by calculating the
Eliassen–Palm flux (hereafter EP flux) divergence, here scaled to units of
zonal acceleration (Andrews et al., 1987):
divEP=∇⋅Fρ0acosϕ,
where F is the EP flux whose
components are
Fϕ=ρ0acosϕuz‾v′θ′‾θz‾-v′u′‾,Fz=ρ0acosϕf-u¯cosϕϕacosϕv′θ′‾θz‾-w′u′‾.f is the Coriolis parameter, a is the Earth's radius, θ is the
potential temperature, ρ0 is the density profile of the atmosphere
and u,v,w are the three-dimensional velocity components in
spherical coordinates λ,ϕ,z, where z is the
log pressure. Overbars indicate zonal mean and primes denote departure from the zonal mean. As shown by Edmon et al. (1980), the EP flux
constitutes a measure of the Rossby wave propagation from one height (z) and
latitude (ϕ) to another, and its divergence (divEP) gives information
about the forcing of the mean flow by the eddies.
Figure 7 displays the monthly evolution in winter of the EP flux and its
divergence for the preindustrial conditions experiment. This analysis shows
that, throughout winter, the wave activity penetrates the stratosphere (as
indicated by the vectors) near 55∘ N, propagates upward and tends
to be increasingly refracted toward the Equator with height. The dissipation
of planetary waves exerts a westward-momentum forcing on the mean flow
between 30 and 70∘ N (as diagnosed by the
Eliassen–Palm flux convergence), which maximizes along the equatorward flank
of the polar night jet where planetary wave breaking is large. This
contributes to eroding and weakening the polar vortex and to a warming of the polar
stratosphere and drives a persistent poleward mass transport in order to
conserve the angular momentum. By mass continuity, this induces an upward
transport at low latitudes and an extratropical downwelling (hence driving
the Brewer–Dobson circulation). Under preindustrial climate conditions, we
note that the wave activity and its interaction with the mean flow peaks in
December or January (Fig. 7c, d) but is already large in November (Fig. 7b).
Therefore, this contributes to slow down the radiatively driven development
of the polar vortex in early winter.
Monthly evolution (October to March) of the Eliassen–Palm flux
(vectors) and its divergence (contours, in m s -1 d-1) under preindustrial
conditions in the Northern Hemisphere.
As shown in Fig. 8, under Eocene conditions, it appears that the planetary
wave activity penetrating the stratosphere in early winter (i.e.,
November–December; Fig. 8b, c) is significantly reduced and deflected
equatorward as revealed by the downward and equatorward pointing of the
EP flux vector in the lower midlatitude stratosphere. This is associated
with an anomalous positive EP flux divergence (i.e., a reduced convergence)
throughout the depth of the stratospheric polar night jet (near
60∘ N), which indicates a substantially reduced westward-momentum
forcing by planetary waves and hence allows a stronger development of the
polar vortex in early winter in comparison with preindustrial conditions. In
contrast, from January (Fig. 8d), the planetary wave activity becomes
significantly larger under Eocene conditions, the westward forcing appears
to be strongly amplified in the upper stratosphere and to progressively
propagate downward in February (Fig. 8e). This is consistent with the
reversal of the zonal-mean zonal wind anomaly in the upper stratosphere, but
also with the overall extremely rapid deceleration of the polar-vortex
strength noted previously (see Figs. 6 and S2). In addition, we
analyzed the residual mass circulation (not shown) derived from the
transformed Eulerian-mean formalism (Andrews et al., 1987). While no clear
changes in the strength of the Brewer–Dobson circulation are found in early
winter between Eocene and preindustrial conditions, late winter
(February–March) reveals an important acceleration in Eocene conditions,
which is consistent with the much stronger wave forcing found throughout the
extratropical stratosphere (Fig. 8e, f). These results are consistent with
a net acceleration of the Brewer–Dobson under Eocene conditions in
comparison with preindustrial conditions as revealed by the younger age of
air (Fig. 2c). Note that the Brewer–Dobson acceleration appears to be more
pronounced in the Northern Hemisphere, where the mean flow and wave activity
anomalies are found to be stronger than in the Southern Hemisphere (not
shown).
Monthly evolution (October to March) of the differences between
the Eocene and preindustrial conditions of the Eliassen–Palm flux and its
divergence. Dotted regions indicate that anomalies are insignificantly
different at the 5 % confidence level according to a t test. Preindustrial
climatology is shown with dashed contours.
Although the large changes in the background state stratospheric circulation
and its seasonal evolution under Eocene conditions in comparison with
preindustrial climate conditions appear to be largely wave-driven, the
origin of the identified changes in the planetary wave activity and its
interaction with the mean flow remains to be determined. Note that this does
not uniquely depend on changes in tropospheric wave sources but also on
changes in the background flow itself, which modulates the wave propagation
and the nature of wave-mean flow interactions. The planetary wave activity
entering the stratosphere can be altered by numerous factors such as
sea surface temperature changes (e.g., Hu et al., 2014), sea-ice changes
(e.g., Kim et al., 2014), wind changes near the tropopause (e.g., Shepherd and
McLandress, 2011; Karpechko and Manzini, 2017) or topography (Shi et al.,
2014). Additional simulations of the Eocene and the preindustrial period with the
atmospheric model (LMDz) without interactive chemistry and with a flat
topography reveal that changes in the topography have first-order effects
on planetary wave activity and hence on the stratospheric dynamics (not
shown). Between the Eocene and the preindustrial conditions, beside large
changes in the topography, important changes in air–sea thermal contrasts,
sea-ice cover and sea surface temperature could all have a substantial
influence on stratospheric circulation. The complexness of these effects and
their possible interactions make an unambiguous attribution impossible in
the absence of a dedicated experimental protocol that is out of the scope for
the present study.
Climate impact of an interactive stratospheric chemistry
Model results shown in the previous section suggest that stratospheric
dynamics and composition were very significantly altered under Eocene hot
conditions in comparison with preindustrial climate conditions. In turn,
these stratospheric changes may also have influenced the establishment of Eocene climate. Nowack et al. (2015) have shown that stratospheric
changes driven by ozone changes can have an impact on the climate
sensitivity in the context of high GHG concentrations for present-day
conditions. The importance of this stratospheric ozone–climate feedback has,
however, not been assessed in the context of Eocene hot climate. In this
section, we estimate the role of this feedback on the overall Eocene climate
response by comparing the EOCENE experiment (i.e., where ozone is calculated
interactively and, hence, is physically consistent with Eocene conditions)
with Eocene_OzRoyer and Eocene_Oz1855
experiments (i.e., where preindustrial ozone climatologies are prescribed in
Eocene simulations). The latter simulations follow the protocol usually
recommended for simulating paleoclimates (e.g., Kageyama et al., 2017).
Table 2 shows the total ozone and temperature changes as well as the
effective radiative forcings induced by the use of an interactive
stratospheric chemistry instead of seasonally varying prescribed
climatologies. All the results in this section are discussed in terms of
25-year means. The effective radiative forcing is computed as the difference
of net radiative flux at the top of atmosphere (TOA) between two atmospheric
simulations (with ozone calculated interactively – with ozone climatology)
as defined in Fig. 8.1.d of Myhre et al. (2013).
Global change in total column ozone, temperature and effective
radiative forcings induced by the use of an interactive stratospheric
chemistry instead of climatologies.
Interactive O3 vs. RoyerInteractive O3 vs. an 11-year mean(EOCENE- EOCENE_OzRoyer)climatology centered on 1855(EOCENE- EOCENE_Oz1855)Change in globally averaged TCO (DU)45.534.2Effective radiative forcing (W m-2)1.41.7Global temperature change (K)0.40.3Stratospheric temperature change (K) above 230 hPa1.41.0
Figure 9 shows the distribution of total column ozone as a function of
latitude for the different configurations. The preindustrial ozone
distribution computed by REPROBUS and the Szopa et al. (2013) ozone
climatology are represented as well. Comparing only the different
preindustrial ozone distributions, the TCO in the experiment with the interactive calculation of ozone (PREIND in black) is higher than those of
the climatologies (OzRoyer in blue, Oz1855 in brown). In addition, the
interactive calculation of Eocene ozone (EOCENE in red) leads to much higher
TCO than in the preindustrial climatologies (blue, brown), the 2000
climatology (green) and the preindustrial interactive ozone simulation
(black). The global mean TCO is increased by about 45 or 35 DU with
respect to the preindustrial OzRoyer or Oz1855 climatologies, respectively.
For the sake of comparison, the global mean TCO had decreased by about 13 DU
only between the 1960s and the end of the 1990s because of the past emissions of
anthropogenic halogenated compounds, and the expected TCO increase at the
2100 horizon is projected to be between 13 and 32 DU depending on future
anthropogenic emissions of GHGs (Bekki et al., 2013; Szopa et al., 2013).
Taking the ozone distribution calculated in the EOCENE simulation as the
reference, the TCO bias in an inappropriate ozone climatology can be 2 times
higher (in the case of the OzRoyer climatology) than the TCO change
calculated by the model between Eocene and preindustrial conditions (EOCENE
versus PREIND; see Sect. 3).
Latitudinal distribution of ozone considered by the circulation
model LMDz when using climatologies from Royer (blue), from Szopa et al. (2013) centered on the year 2000 (green) or centered on the year 1855
(maroon) or interactively computed by REPROBUS for Eocene conditions (red)
or preindustrial conditions (black).
The TCO difference between the EOCENE ozone distribution and ozone
climatologies peak at midlatitudes (about 40∘), reaching almost
70 DU for the Oz1855 climatology and about 100 DU for the OzRoyer
climatology (Fig. 10a). In order to identify the regions responsible
for the general increase in TCO calculated from preindustrial to Eocene
conditions, the zonal mean distributions of ozone difference between EOCENE
and Oz1855 climatologies are plotted in DU km-1 in Fig. 10b.
The TCO increase in EOCENE is largely due to an enhancement in ozone in the
upper stratosphere. The TCO change in the tropics is very moderate because
the upper-stratospheric ozone enhancements are more or less compensated for by lower-stratospheric ozone decreases brought about by the acceleration of the
Brewer–Dobson circulation, namely the faster ascent in the tropics (Sect. 3).
TCO enhancements peak at midlatitudes because the ozone concentration
increases reach down to 150 hPa at midlatitudes, again certainly linked
to the acceleration of the Brewer–Dobson circulation and more specifically
the faster descent at mid- and high latitudes. Below 200 hPa, around the
tropopause region, extratropical EOCENE ozone concentrations are lower than
in the Oz1855 climatology, mostly because of the rise in the tropopause height
(Sect. 3).
Total column ozone (TCO) changes (in Dobson units or DU) between
the EOCENE simulation and the EOCENE-Oz1855 simulation (considering the 1855
ozone climatology) (a) and TCO changes between the EOCENE simulation
and the two climate-only simulations considering the (solid) 1855 ozone
climatology and (dashed) Royer ozone parameterization (b).
Ozone changes naturally lead to temperature changes, especially in the
stratosphere where ozone and temperature are closely coupled. Figure 11
shows the zonal mean distribution of temperature difference between EOCENE
and EOCENE_Oz1855 simulations. The impact is weak below 400 hPa since SSTs are fixed and identical in all the Eocene simulations
(EOCENE, EOCENE_OzRoyer, EOCENE_Oz1855). The
change in zonal mean temperatures below 400 hPa does not exceed 0.15 K but
can almost reach 0.5 K for the northern polar latitudes when the interactive
ozone simulation (EOCENE) is compared to the EOCENE_OzRoyer
(not shown). In contrast, temperatures above about 200 hPa are significantly
impacted by the choice of ozone distribution used in the model. Temperatures
are more than 2.5 to 3 K higher at middle and high latitudes in both
hemispheres when ozone is calculated interactively instead of using the
preindustrial 1855 climatology. The largest differences are found near the
stratopause region (above 5 hPa), in the lower polar lower
stratosphere(∼130 hPa) and in the middle tropical
stratosphere (∼60 hPa), where temperatures are, respectively, higher than 6 K, higher than 4 K and lower than 3.5 K in the simulation with
interactive ozone compared to the one with the preindustrial climatology.
Stratospheric temperature changes (K) between the EOCENE
simulation and the climate-only simulation with the 1855 ozone climatology.
Do we need to consider stratospheric ozone feedback in deep-past
simulations?Impact on climate
The consideration of a stratospheric ozone compatible with the Eocene
conditions perturbs the radiative balance compared to the use of a
preindustrial climatology. The net radiative change (shortwave + longwave)
between the simulation with interactive chemistry and the simulation with
preindustrial climatology corresponds to a 1.7 W m-2 effective
radiative forcing (RF). This radiative forcing results from combined
positive RFs in the longwave (LW) and shortwave (SW) simulated in the
tropics. Beyond 50∘ (north and south), the positive SW RF is partly counterbalanced
by a negative longwave RF (see Table 3 and Fig. 12). This radiative
forcing from the stratospheric ozone response represents a positive climate
feedback, which is commonly ignored in Eocene climate simulations. In order
to estimate the potential impact of an interactive ozone on surface
temperature under Eocene conditions, we consider a large range of
climate sensitivity from 0.4 to 1.2 K (W m-2)-1 (Knutti et al.,
2017). Given such a broad range, the surface temperature response to this
stratospheric forcing could range from 0.7 to 2.0 K (assuming an effective
radiative forcing of 1.7 W m-2). The surface temperature response to a
specific radiative forcing depends on the considered climate conditions and
on the nature of the climate forcer. Considering an interactive
stratospheric ozone chemistry under a 4×CO2 climate perturbation, Nowack
et al. (2015) found a climate sensitivity of 1.05 K (W m-2)-1 in
their ESM. Applying this climate sensitivity, the surface temperature change
associated with the ozone feedback would be about 1.8 K when considering
interactive stratospheric chemistry (compared with the EOCENE_Oz1855 run). The climate sensitivity to ozone change can obviously vary from
one ESM to another since, for example, the sensitivity of the Brewer–Dobson
circulation to climate is highly model-dependant (SPARC CCMVal, 2010).
Nonetheless, this estimation allows us to discuss the importance of
considering this chemistry–climate feedback when attempting to simulate
greenhouse paleoclimates. According to the IPCC AR5 report (IPCC, 2013), the
global land surface air temperature anomaly is +12.7 K for the early
Eocene climatic optimum (Masson-Delmotte et al., 2013). This estimation is
based on simulations from several models analyzed by Lunt et al. (2012), for which there was no common modeling protocol (e.g., CO2 being in the
2×[CO2]PI to 16×[CO2]PI range). One of these models, the HadCM model, estimates that
the effect of changing non-CO2 boundary conditions (topography,
bathymetry, solar constant and vegetation) for Eocene conditions leads to a
1.8 K increase in the global mean surface air temperature (to be compared to
a 3.3 K increase when doubling the CO2). The feedback of stratospheric
ozone on surface air temperature could thus represent about 15 % of the
total temperature anomaly reported between Eocene and preindustrial
conditions and be as important as the effect of external forcings.
Differences in shortwave and longwave radiative fluxes at different
vertical levels between the EOCENE simulation and the climate-only
simulation with the 1855 ozone climatology.
Differences in radiative fluxes as a function of latitude between
the EOCENE simulation and the climate-only simulation with the 1855 ozone
climatology.
In strong greenhouse climate, the terrestrial carbon and nitrogen cycles are
intensified, releasing high CH4 and N2O in the atmosphere
(Beerling et al., 2011). The effect of changing N2O and CH4 in the
troposphere (including the H2O increase in the stratosphere induced by
the CH4 increase) has been assessed by Beerling et al. (2011). These
authors find, with the STOCHEM model, a 2.1 K increase in global surface
temperature due solely to the tropospheric composition changes. Our results
suggest that the effect of the stratospheric ozone feedback on surface
temperatures is of similar importance.
Impact on tropospheric conditions
Using an ESM including chemistry, Unger and Yue (2014) found that under
warm and high methane Pliocene conditions, the stratospheric ozone burden
was 5 % higher than the preindustrial one. This stratospheric ozone
increase resulted in a 20 % reduction in the tropospheric photolysis rate
of ozone (O3+hv->O1D+O2) that leads to
the formation of OH, the hydroxyl radical. This radical is the main oxidant
of the troposphere and its decrease (of 20 % to 25 % in the Unger and Yue
simulations) impacts the lifetime of chemical species and in particular
CH4. Our simulations show a 7.2 % increase in the stratospheric ozone
burden when comparing the EOCENE and the PREIND simulation (both including
interactive chemistry) and an 8.8 % difference when comparing the EOCENE
simulation to the EOCENE_Oz1855.
In addition, we estimate the change in surface UV radiations and ozone
photolysis in the Eocene conditions. Using the radiation transfer model
Quick TUV Calculator with a pseudo-spherical discrete ordinate four streams
(http://cprm.acom.ucar.edu/Models/TUV/Interactive_TUV/, last access: 2 June 2019), we estimate the effect of the stratospheric ozone increase on the
ozone photolysis, which controls the OH production. Considering a 65 DU
change at a 50∘ latitude (corresponding to the maximum of Fig. 10), the photolysis rate at the surface decreases by 25 %. A decrease in
the ozone photolysis rate would induce a significant decrease in OH and
hence in the tropospheric oxidizing capacity, thus making CH4 longer
lived and reinforcing its effect on climate. It would also impact the
overall tropospheric chemistry.
Conclusion
The stratospheric dynamics and ozone layer respond to – and interact with –
atmospheric variations (climate, tropospheric GHG content). In this study,
we simulate these interactions in the case of the hot Eocene climate using a
chemistry–climate model. We characterize the changes in ozone and middle-atmospheric dynamics-induced hot Eocene climate conditions characterized by
a 4×CO2 climate, elevated concentrations of CH4 and N2O, and
substantial changes in surface boundary conditions (e.g., sea surface
temperature, sea-ice cover or topography) compared to
preindustrial climate. The climate impact of the stratospheric response
under hot conditions is also discussed.
Comparing the Eocene simulation with a preindustrial simulation, we find a
sharp increase in ozone in the upper stratosphere (reaching 40 % at 2–3 hPa
in the tropics) linked to the strong cooling of the stratosphere (up to -12 K at 10 hPa), which slows down the chemical destruction of ozone. Meanwhile,
ozone is greatly reduced in the lower tropical stratosphere (up to 40 %)
due to the intensification of Brewer–Dobson circulation. These results are
in agreement with previous modeling studies that considered current
tropospheric composition and a 4×CO2 climate change.
As a consequence of the opposite ozone changes in the tropics (enhanced
ozone in the upper stratosphere, reduced ozone in the lower stratosphere),
the tropical total column ozone (TCO) is not affected much by the difference
in climate between the Eocene and preindustrial periods. On the contrary, at
midlatitudes and, to a lesser extent, in the polar regions, the TCO is
considerably increased. The TCO meridional distribution is also strongly
modified, exhibiting particularly pronounced midlatitude maxima and
a steeper negative poleward gradient from these maxima. These changes in
meridional distribution reflect significant polar-vortex changes during the
winter or early spring, especially in the Northern Hemisphere. The polar vortex
becomes stronger and more extended equatorward under Eocene conditions,
thus reinforcing the isolation of the polar-vortex air masses from the
midlatitudes in comparison with preindustrial conditions. In our
simulations, the reinforcement of the stratospheric polar vortex under
Eocene conditions and the acceleration of the stratospheric overturning
circulation (which seems contradictory at first) is consistent with a
reduced intensity of the planetary wave activity and its interaction with
the mean flow in early winter and, conversely, a strongly amplified wave
activity and interaction with the mean flow in late winter compared to
preindustrial conditions.
We then explore the possible role of the stratospheric ozone response in the
establishment of Eocene climate. For that purpose, we compare the
simulations with interactive ozone with simulations forced by the use of
preindustrial ozone climatologies. The difference in global mean TCO between
the Eocene simulation and simulations using preindustrial climatologies is
2 to 3 times higher than the change in ozone observed between 1960 and
end of the 1990s (the minimum TCO) and of the same order as the changes
projected between 2000 and 2100. The ozone increase in the upper
stratosphere in the case of Eocene interactive ozone warms the atmosphere by
up to 3 K above 230 hPa. In the tropical lower stratosphere, zonal mean
temperatures are up to 3.5 K lower for the Eocene stratospheric ozone
compared to the preindustrial ozone. These changes in the thermal structure
of the middle atmosphere could, via atmospheric circulation teleconnections,
have significant regional consequences. Using the sensitivity of surface
temperatures to stratospheric ozone changes determined by Nowack et al. (2015)
(though climate sensitivity varies among climate–chemistry models), we
estimate the contribution of stratospheric ozone feedback to surface
temperature change in Eocene hot climate simulations. We find that it is
potentially as important as the effects of non-CO2 boundary conditions
(topography, bathymetry, solar constant and vegetation) or uncertainties
due to gaseous tropospheric chemistry. The results suggest that future
studies exploring long-standing Cenozoic warm-climate questions, such as
the varying latitudinal temperature gradient during hothouse periods, would
benefit from exploring and integrating – even if costly in terms of computing time – the
feedbacks of ozone on atmospheric temperatures rather than prescribing
preindustrial values.
Data availability
The main model outputs are publicly accessible at
10.14768/201906141.01 (Szopa et al., 2019).
The supplement related to this article is available online at: https://doi.org/10.5194/cp-15-1187-2019-supplement.
Author contributions
The idea of the study, the design of numerical simulations and the
radiative forcing analysis come from SS. RT performed
all the analysis on atmospheric dynamics. SvB provided the Eocene
boundary conditions. SS and RT prepared the first
draft of the paper. All coauthors contributed to its editing.
Competing interests
The authors declare that they have no conflict of
interest.
Acknowledgements
The authors are
thankful to the NCAR Atmospheric Chemistry Division (ACD) for the
distribution of the NCAR/ACD TUV: Tropospheric Ultraviolet & Visible
Radiation Model (http://cprm.acom.ucar.edu/Models/TUV/Interactive_TUV/) and the
availability of their quicktool. We thank Yannick Donnadieu for preparing the
paleogeographical conditions for Eocene simulations and Marion Marchand for
her advice in the use of the REPROBUS model and Jean-Louis Dufresne for
helpful discussion on radiative forcing.
Financial support
This research has been supported by the Agence Nationale de la Recherche (PALEOx project, grant no. ANR-16-CE31-0010). This work was granted access to the HPC resources of TGCC under the allocation 2017-A0050102212 made 40 by GENCI (Grand Equipement National de Calcul Intensif).
Review statement
This paper was edited by Arne Winguth and reviewed by two anonymous referees.
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