An accurate quantification of the stratospheric ozone feedback in climate change simulations requires knowledge of the ozone response to increased greenhouse gases. Here, an analysis is presented of the ozone layer response to an abrupt quadrupling of CO 2 concentrations in four chemistry–climate models. The authors show that increased CO 2 levels lead to a decrease in ozone concentrations in the tropical lower stratosphere, and an increase over the high latitudes and throughout the upper stratosphere. This pattern is robust across all models examined here, although important intermodel differences in the magnitude of the response are found. As a result of the cancellation between the upper- and lower-stratospheric ozone, the total column ozone response in the tropics is small, and appears to be model dependent. A substantial portion of the spread in the tropical column ozone is tied to intermodel spread in upwelling. The high-latitude ozone response is strongly seasonally dependent, and shows increases peaking in late winter and spring of each hemisphere, with prominent longitudinal asymmetries. The range of ozone responses to CO 2 reported in this paper has the potential to induce significant radiative and dynamical effects on the simulated climate. Hence, these results highlight the need of using an ozone dataset consistent with CO 2 forcing in models involved in climate sensitivity studies.

The present paper documents the ozone responses to CO 2 obtained in the different CCMs. The ozone responses in the four models will then be used in a follow-up study to quantify the feedback in the form of radiative forcing, and dynamical effects of ozone and its zonal asymmetries on the atmospheric circulation.

We examine the ozone response to an abrupt quadrupling of CO 2 in four different CCMs. Using four different models allows us to identify the robust features, and to quantify the intermodel spread. CO 2 is the only external forcing in these runs: this facilitates the attribution of the forced response. Moreover, the large instantaneous forcing from a quadrupling of CO 2 concentrations allows us to distinguish fast and slow responses ( Gregory and Webb 2008 ; Taylor et al. 2012 ), thus providing insights into the mechanisms driving the ozone response. Last, the longitudinal structure of the ozone response is analyzed in detail to highlight asymmetries in the ozone response, a feature that is presently omitted in ozone forcing datasets ( Cionni et al. 2011 ).

There is robust modeling evidence suggesting that anthropogenic greenhouse gases (GHGs), via their influences on stratospheric temperature and the Brewer–Dobson circulation (BDC), will greatly modify the future distribution of ozone in the stratosphere ( WMO 2014 , chapter 2.4.2). More specifically, GHGs induce stratospheric cooling, but also strengthen the BDC. The cooling and BDC strengthening have opposite influences on the ozone layer in the tropics: radiative cooling slows down ozone catalytic cycles and affects gas-phase ozone photochemistry (thus increasing ozone concentrations), while the strengthening of the BDC enhances advection of ozone-poor air in the tropical lower stratosphere, thus decreasing ozone concentrations ( Shepherd 2008 ). However, the exact contribution of single forcing agents is unclear.

An accurate quantification of the effects of anthropogenic emissions on the ozone layer is a key step toward making accurate predictions of the future ozone evolution. Assessing the ozone response to anthropogenic forcings is also a step toward improved understanding of the coupling between atmospheric composition and climate ( Isaksen et al. 2009 ).

All four models have model tops well above 1 hPa (~50 km) and have a well-resolved stratosphere. Therefore, they are considered “high top” models ( Charlton-Perez et al. 2013 ). Most importantly, they include fully interactive stratospheric ozone chemistry: thus, the interplay between ozone chemistry, radiation, and dynamics is fully represented in all of them. There are some differences in tropospheric ozone chemistry, due to the representation of feedbacks between climate and lightning NOx. In GISS-E2-H, GFDL CM3, and CESM(WACCM), lightning NOx sources are interactive and thus respond to changes in climate, while in SOCOL, they are prescribed through a climatological source of 4 Tg (N)/yr −1 . The complexity of the tropospheric chemistry mechanism differs among models, with some (e.g., GFDL CM3) including more reactions and species than others [SOCOL and CESM(WACCM)]. However, ozone responses in the troposphere are dwarfed by those in the stratosphere, as shown below.

The SOCOL model has a spectral resolution of T42, corresponding to 2.8° longitude by 2.8° latitude, 39 vertical levels, and a top at 0.01 hPa (~80 km). Ocean coupling is provided by the ocean–sea ice model Max Planck Institute Ocean Model. An accurate description of the model physics and chemistry is given in Stenke et al. (2013) . Atmospheric chemistry is calculated through 140 gas-phase reactions, 16 heterogeneous reactions, and advection of 41 chemical species. The transport of the chemical species, including ozone, is calculated by the advection scheme of the middle-atmosphere ECHAM5.

The CESM(WACCM) model has a resolution of 1.9° longitude by 2° latitude and 66 vertical layers, with a model top at 5.96 × 10 −6 hPa (~140 km). The ocean component is provided by the Parallel Ocean Program, version 2 (POP2). CESM(WACCM) is fully documented in Marsh et al. (2013) . The model includes a fully interactive stratospheric chemistry module, based on version 3 of the Model for Ozone and Related Chemical Tracers (MOZART; Kinnison et al. 2007 ), which involves 217 gas-phase reactions, and the advection of a total of 59 species. This version of CESM(WACCM) also includes a simplified representation of tropospheric chemistry, which is limited to methane and CO oxidation [see Marsh et al. (2013) for more details]. We note that CESM(WACCM) does not include aerosol indirect effects.

The GFDL CM3 model has a resolution of 2.5° longitude by 2° latitude and 48 vertical layers, with a model top at 0.017 hPa (~60 km). The ocean model component of CM3 is the Modular Ocean Model (MOMp1; Griffies et al. 2005 ). As in GISS-E2-H, this model includes clouds–aerosol interactions. The atmospheric component includes modules for tropospheric and stratospheric chemistry, based on Horowitz et al. (2003) and Austin and Wilson (2006) , respectively. Tropospheric and stratospheric chemistry modules have been merged, which implies extending the tropospheric chemistry module to include 63 chemical species, halogens, atomic hydrogen, and oxygenated species, as well as PAN and other ozone precursors. Details of the GFDL CM3 model physics can be found in Donner et al. (2011) .

The GISS-E2-H model has a resolution of 2.5° longitude by 2° latitude and 40 vertical layers, with a model top at 0.1 hPa (~60 km), and is coupled to the Hybrid Coordinate Ocean Model (HYCOM). The model includes the first aerosol indirect effect (i.e., the impact of aerosols on cloud microphysical processes). It employs 51 species for gas-phase chemistry interacting via 156 reactions. Ozone is prognostic both in the stratosphere and in the troposphere and thus evolves with the atmospheric state ( Shindell et al. 2013 ). Tropospheric chemistry includes basic NOx, HOx, Ox, and CO-CH 4 chemistry as well as peroxyacyl nitrates (PANs) and hydrocarbons. This configuration is commonly referred to as the Tracers of Chemistry, Aerosols, and their Direct and Indirect Effects (“TCADI”) and is identified as p3 (physics-version = 3) in the CMIP5 archive. More details about the model physics and dynamics are given in Schmidt et al. (2014) .

For our analysis, we employ four atmosphere/ocean coupled chemistry–climate models: the Goddard Institute for Space Studies Model E2-H (GISS-E2-H), the Geophysical Fluid Dynamics Laboratory Climate Model, version 3 (GFDL CM3), the Community Earth System Model (Whole Atmosphere Community Climate Model), version 4 [CESM(WACCM)], and the coupled model for studies of Solar-Climate-Ozone Links, version 3 (SOCOL).

For two of the models (i.e., GISS-E2-H and GFDL CM3), we use the data available on the CMIP5 archive. For CESM(WACCM), we use the same data analyzed in Marsh et al. (2016) and Chiodo and Polvani (2017) . For SOCOL, we analyze the output documented in Muthers et al. (2014) . Where it is shown, we assess the equilibrium response of ozone to CO 2 by taking differences between the climatology obtained from the last 50 years of the 4×CO 2 integrations and the climatologies obtained from the 150 yr-long PI control integrations. After 100 years, ozone trends are found to be very small. Thus, these climatological differences will be referred to as “equilibrium response,” although they do not strictly represent a new steady state.

In both control and 4×CO 2 runs, ODSs and ozone precursors are kept at PI levels. This implies that any changes in polar stratospheric clouds formation (e.g., due to CO 2 -induced stratospheric temperature changes) will not have a sizable effect on stratospheric ozone. Imposing a CO 2 forcing on an atmosphere with “present day” levels of ODSs could have an effect on heterogeneous chemistry, but would be inconsistent with the approach employed in CMIP5 studies to assess forcing, feedbacks, and climate sensitivity ( Andrews et al. 2012 ).

We analyze two different forcing scenarios from each of the CCMs: a preindustrial (PI) control and an abrupt 4×CO 2 scenario of equal length (150 yrs long), in which atmospheric CO 2 is instantaneously quadrupled at the beginning of the run. It is important to stress that ODSs and tropospheric ozone precursor emissions are held fixed to PI levels in both integrations: this is a key distinction between 4×CO 2 forcing and the emission scenarios analyzed in earlier studies (e.g., Oman et al. 2010 ; Eyring et al. 2010 , 2013 ; Iglesias-Suarez et al. 2016 ).

There is some coherence between intermodel spread in tropical stratospheric ozone and temperature. For example, SOCOL shows the largest ozone decrease at 30 hPa, and is also the model with the largest cooling in response to CO 2 , between 50 and 10 hPa ( Fig. 3 ). The opposite is seen in CESM(WACCM): a weaker TLS ozone decrease in this model could explain the weaker cooling at 30–10 hPa. This suggests that ozone responses may contribute to intermodel spread in the stratospheric cooling because of increased CO 2 levels. Nevertheless, there is no relationship between temperature and ozone response in GISS-E2-H, suggesting that other processes, perhaps dynamical cooling or stratospheric water vapor (e.g., due to intermodel differences in the strength of the stratospheric water vapor feedback; see Dessler et al. 2013 ), may also contribute to the intermodel spread in the stratospheric temperature response to CO 2 .

To bring out the intermodel differences in the tropical ozone response to CO 2 , we show the annual-mean tropical average (30°S–30°N) profile of ozone mixing ratios in Fig. 4 . First, we note differences in the location of the peak in the upper-stratosphere (3–5 hPa) ozone increase, with GISS-E2-H and CESM(WACCM) showing a peak at higher altitudes than SOCOL. Second, while models agree in the location of the maximum ozone decrease at 30 hPa, there is significant intermodel spread in amplitude; the ozone decrease ranges between 0.2 ppmv [GISS-E2-H and CESM(WACCM)] and 1.0 ppmv (SOCOL). Third, one can easily see that tropospheric ozone changes are extremely small compared to those occurring in the stratosphere. In the following section, we will show that the spread in tropical lower-stratospheric ozone is consistent with intermodel differences in the BDC, and tropospheric temperature.

There are also some notable intermodel differences in the magnitude of the stratospheric ozone response in the tropics. In the upper stratosphere, the ozone increase ranges from 40% in CESM(WACCM) and GISS-E2-H, to 30% in SOCOL and GFDL CM3. In the TLS, the decrease in ozone concentrations ranges from 50% in SOCOL to 30% in CESM(WACCM). These intermodel differences are more evident when looking at ozone volume mixing ratio ( Fig. S1 ). Some differences among models are also present in their PI control climatology ( Fig. S2 ), although these are generally smaller than the response to CO 2 , especially at low latitudes.

The SOCOL model is consistent with the other models in projecting an ozone increase in the tropical and subtropical upper troposphere (300 hPa), despite the lacking response in lightning NOx emissions to CO 2 increase in this model. This suggests that tropospheric ozone increases can be driven by other processes, such as stratosphere–trosphere exchange (STE; Hegglin and Shepherd 2009 ; Garny et al. 2011 ). The specific pattern, with a positive ozone response extending from the subtropical upper troposphere poleward and upward to the lower stratosphere in the midlatitudes, is a further indication that STE could contribute to the tropospheric ozone response to CO 2 .

In the troposphere, a dipole of ozone increases in the midtroposphere and decreases close to the tropopause layer is seen in all models. The pattern of tropospheric ozone response to CO 2 has been linked to enhanced NOx lightning, and uplifting of the tropopause (i.e., ozone-poor tropospheric air replacing stratospheric air; Dietmüller et al. 2014 ). In the middle troposphere, enhanced NOx lightning can result from changes in both the intensity (depth) of individual convective events, and the overall frequency of convection with warming ( Banerjee et al. 2014 ). Enhanced NOx in the free troposphere can lead to more efficient ozone production via cycling of HOx and NOx radicals ( Brasseur and Solomon 2005 ).

In the lower stratosphere, the decrease in ozone concentrations is likely due to an acceleration of the BDC ( Butchart 2014 ); both stratspheric cooling and the BDC strengthening are robust features in climate change simulations, and also dominate the ozone response to 4×CO 2 .

The upper-stratospheric ozone increase has been understood to be a consequence of changes in odd oxygen loss cycles because of CO 2 -induced cooling ( Haigh and Pyle 1982 ; Jonsson et al. 2004 ). In this region, all models show a similar cooling of up to 16 K ( Fig. 3 ). Assuming photochemical equilibrium, and following the analytical calculation presented in Jonsson et al. [2004 ; their Eq. (7)], a −16 K temperature change at 1–5 hPa would lead to an 11% increase in the reaction rate coefficient involved in recombination (O + O 2 + M → O 3 ), and a 44% decrease in the reaction rate coefficient involved in ozone destruction (O 3 + O → 2O 2 ). Combining the effect of both reaction rate coefficients, and assuming no changes in OH, NO 2 , and ClO concentrations, we calculate an ozone increase of ~27% at 5 hPa, which is close to the values calculated by the models and explains the robustness of the upper-stratospheric ozone signal in the different CCMs.

Relative annual-mean zonal-mean ozone response in (a) CESM(WACCM), (b) GFDL CM3, (c) GISS-E2-H, and (d) SOCOL (units: %). The thick violet solid (stippled) line identifies the tropopause in each of the models for the control (4×CO 2 ) experiment, calculated using the WMO lapse rate definition. Regions that are not stippled are statistically significant (at the 99% level), according to the t test.

Relative annual-mean zonal-mean ozone response in (a) CESM(WACCM), (b) GFDL CM3, (c) GISS-E2-H, and (d) SOCOL (units: %). The thick violet solid (stippled) line identifies the tropopause in each of the models for the control (4×CO 2 ) experiment, calculated using the WMO lapse rate definition. Regions that are not stippled are statistically significant (at the 99% level), according to the t test.

The equilibrium response in zonal-mean ozone, calculated as relative change, along with the tropopause diagnosed using the WMO definition ( WMO 1992 ), 1 is plotted in Fig. 2 . In the stratosphere, we identify a robust pattern of ozone response in the low latitudes, which consists of an increase by up to 30%–40% in the upper stratosphere (1–10 hPa), and a decrease of similar magnitude in ozone in the tropical lower stratosphere (TLS) (30–100 hPa). Relative changes near the tropopause are large (30%–50%). However, in (absolute) mixing ratio terms, the decreases in the lower stratosphere are smaller than the increases in the upper stratosphere (see Fig. S1 ). Despite their small size in terms of volume mixing ratio, ozone changes in the lower stratosphere are particularly important for the global energy budget ( Lacis et al. 1990 ).

The time evolution of the global mean surface temperature response to 4×CO 2 in the four models is shown in Fig. 1 . All models exhibit rapid surface temperature increase over the first 10–20 years following the CO 2 quadrupling, and then warm at a smaller and more model-dependent rate. Over the simulated period, the warming ranges between 4.2 K (GISS-E2-H) and 5.8 K (SOCOL). Over the first 150 years, the warming in CMIP5 models in CO 2 quadrupling experiments typically ranges between 3.0 K and 6.2 K [see Table S1 in Grise and Polvani (2014) ]. The key point here is that the four CCMs span over a good fraction (~50%) of the existing spread in climate sensitivity (measured as surface temperature response to 4×CO 2 ) across the CMIP5 models.

Next, we vertically integrate the response displayed in Fig. 2 to quantify the equilibrium response in total column ozone. First, we integrate over the whole column to yield the total column ozone in Dobson units (DU) (named hereafter “TO3”). Then, we repeat the integration for the troposphere only (“TRO3”). In the stratosphere, the existence of opposite responses (see Fig. 4) motivates separating two distinct regions: the lower stratosphere, defined as the atmospheric layer between the tropopause and 20 hPa (“LSO3”), and upper stratosphere, defined as the layer between 20 hPa and 1 hPa (“USO3”). Figure 5 shows the latitudinal structure of the equilibrium response of TO3, TRO3, LSO3, and USO3 (Figs. 5a–d, respectively) to a quadrupling of CO 2 .

Fig . 5. View largeDownload slide Zonal-average column ozone response to 4×CO 2 ; (a) total, (b) tropospheric, (c) lower-stratosphere, and (d) upper-stratosphere partial ozone column. The lower stratosphere is defined as the atmospheric layer between the tropopause and 20 hPa, while the upper stratosphere is defined as the layer between 20 hPa and 1 hPa (units: DU). Error bars span over the 2σ uncertainty, represented by the standard error of the mean. Fig . 5. View largeDownload slide Zonal-average column ozone response to 4×CO 2 ; (a) total, (b) tropospheric, (c) lower-stratosphere, and (d) upper-stratosphere partial ozone column. The lower stratosphere is defined as the atmospheric layer between the tropopause and 20 hPa, while the upper stratosphere is defined as the layer between 20 hPa and 1 hPa (units: DU). Error bars span over the 2σ uncertainty, represented by the standard error of the mean.

Starting from Fig. 5a, we see that all models project a total column ozone increase at high latitudes, with a larger increase in the NH than in the SH (Fig. 5a). On the other hand, tropical column ozone responses are small. This pattern is consistent with the response in the most extreme RCP8.5 scenario (cf. Butler et al. 2016, Fig. 1 therein), despite the very different forcings employed here. Most importantly, the stratospheric ozone response is the dominant contributor to the latitudinal pattern of TO3 (Figs. 5c,d). Further, we can see a large cancellation between USO3 increases (Fig. 5c) and LSO3 decreases (Fig. 5d), resulting in a small TO3 response in the tropics (Fig. 5a). The tropospheric column ozone response is generally small (less than 5 DU), which is possibly due to cancellations between ozone increase in the middle troposphere, and decrease near the tropopause in Fig. 2. In USO3, all models show a similar increase of 20 DU, with the exception of SOCOL, which shows larger values (30–35 DU) because of the lower altitude of the upper-stratospheric peak in Fig. 4 (and hence larger effect on ozone number density).

We also note a significant intermodel spread in the magnitude of high-latitude ozone increase, and in the sign of the response in tropical ozone column: this spread is almost entirely generated in the LSO3 (Fig. 5d). At high latitudes, the ozone increase is largest in GISS-E2-H (50 DU), and smallest in SOCOL (10–20 DU). In the tropics, the models with the largest LSO3 decrease also exhibit a TO3 decrease; this is the case for SOCOL and GFDL CM3. This suggests that the uncertainty in the sign of the tropical TO3 response (Fig. 5a) is mostly due to uncertainty in the magnitude of the LSO3.

It is widely believed that the projected changes in LSO3 are due to the acceleration of the BDC over the twenty-first century (Butchart 2014). Thus, a possible source of spread in the tropical ozone is stratospheric upwelling. Ideally the BDC would be diagnosed using the transformed Eulerian-mean (TEM) winds (Andrews et al. 1987). Here, we calculate upwelling at the 100-hPa level, as the Eulerian-mean velocity field averaged between turnaround latitudes (22°N–22°S) at the 100-hPa level resembles the TEM residual velocities [see chapter 3 in Andrews et al. (1987)]. Thus, at this level provides an approximate measure of the strength of the upwelling branch of the BDC. The scatterplot of ozone and upwelling responses at 100 hPa is shown in Fig. 6 for total (Fig. 6a) and lower-stratospheric column ozone (Fig. 6b). The negative correlation between changes in upwelling and ozone is highly significant, indicating that models with the largest upwelling response to 4×CO 2 forcing (SOCOL and GFDL CM3) also project the largest decrease in lower-stratospheric column ozone (Fig. 6b), showing the importance of the BDC in determining the ozone response in the TLS. Similar results are obtained using at 70 hPa (not shown). The decrease in lower-stratospheric ozone in SOCOL and GFDL CM3 is sufficiently large to overcompensate the increase in upper-stratospheric ozone (USO3), thus resulting in a negative change in total column ozone (Fig. 5a). We thus conclude that the uncertainty in the sign of the tropical ozone response stems from the intermodel spread in the strengthening of the ascending branch of the BDC.

Fig . 6. View largeDownload slide Scatterplot of upward velocity change at 100 hPa in response to 4×CO 2 and (a) total column ozone, and (b) lower-stratospheric ozone column–averaged in the tropical region (22°S–22°N). Fig . 6. View largeDownload slide Scatterplot of upward velocity change at 100 hPa in response to 4×CO 2 and (a) total column ozone, and (b) lower-stratospheric ozone column–averaged in the tropical region (22°S–22°N).

Interestingly, models with the largest upwelling response, such as SOCOL and GFDL CM3, are also the models with the largest tropical tropospheric warming (Fig. 3). A close relationship between tropospheric warming rates and upwelling is also evident from the transient response in the four models (Fig. S3). This suggests a possible relationship between intermodel spread in stratospheric upwelling, decreased ozone concentrations in the TLS, and climate sensitivity. Decreased ozone in the TLS can exert a substantial radiative forcing (Hansen et al. 2005), which might have important implications for tropospheric climate.

Up to this point, we have looked at the equilibrium response in ozone. But what time scales are needed to reach an equilibrated state? The instantaneous quadrupling of CO 2 is an idealized forcing, which allows a separation of fast and slow responses, and is thus useful to elucidate the mechanisms driving the oppositely signed responses in USO3 and LSO3. Figure 7 shows the time series of the response in tropical averaged USO3 (Fig. 7a) and LSO3 (Fig. 7b). The USO3 increase occurs instantaneously upon quadrupling CO 2 concentrations, while most of the LSO3 decrease takes place over the first 2–3 decades. This behavior clearly hints at very different processes driving the two responses, which are discussed next.

In the upper stratosphere, all models show similar cooling of up to 16 K at 1 hPa (see Fig. 3): this radiatively induced cooling occurs instantaneously upon increasing CO 2 (not shown), changing the reaction rates involved in the Chapman cycle, resulting in increased ozone concentrations (Haigh and Pyle 1982; Jonsson et al. 2004). On the other hand, decreased lower-stratospheric ozone concentrations are associated with enhanced upwelling (Shepherd 2008). It has been suggested that changes in upwelling occur in response to a strengthening of the upper flanks of the subtropical jets, which pushes the critical layers upward, allowing more wave activity to penetrate into the subtropical lower stratosphere (Shepherd and McLandress 2011). The strengthening of the subtropical jets is caused by warming in the upper tropical troposphere, which is in turn a result of changes in convection and thus tropospheric lapse rate. Tropical stratospheric upwelling is tightly coupled with the evolution of upper-tropospheric temperature (Fig. S3). Hence, ozone changes in the TLS proceed at a slower pace than changes in the upper stratosphere, where ozone is mostly in photochemical equilibrium and where the concentrations are governed primarily by (fast) gas-phase reactions that are temperature dependent (Sander et al. 2006).

Another way of splitting fast and slow responses would be to compare ocean-coupled with atmosphere-only simulations using fixed SSTs. Unfortunately, these runs are only available for CESM(WACCM), but not for the other three models. In CESM(WACCM), we find an ozone increase in the upper stratosphere, which closely resembles that observed at 40–50 km in Fig. 2a (not shown). On the other hand, the ozone decrease in the TLS region is about 10% and thus much weaker than in the coupled runs, confirming the role of surface warming and the consequent BDC strengthening in driving the ozone response in this region.

In summary, these results suggest that the tropical ozone response to 4×CO 2 exhibits two different regimes: a fast response in the upper stratosphere, which is radiatively controlled via changes in gas-phase chemistry, and a slower—and opposite—response in the lower stratosphere, where ozone is dynamically controlled. This is consistent with the lifetime of ozone in both regions, which is mostly determined by photochemistry in the upper stratosphere, and transport below 20 hPa (Brasseur and Solomon 2005). Thus, the same processes that determine the background ozone distribution are also key in driving its response to 4×CO 2 .