The Intergovernmental Panel on Climate Change Sixth Assessment Report (IPCC AR6) assessed contributions to observed warming since industrialization. This assessment found that the direct, indirect, and semi-direct effects of aerosols are altogether the largest sources of uncertainty for anthropogenic radiative forcing . Secondary organic aerosols (SOAs) account for a significant portion of aerosols in the troposphere. SOAs are in turn produced through a range of chemical reactions involving volatile organic compounds (VOCs), and the spatial distributions of SOAs and VOCs are coupled to atmospheric dynamics. Thus, understanding the climate effects of tropospheric VOC chemistry is critical to reducing uncertainty in assessing contributors to both past and projected climate change.

VOCs include biogenic VOCs (BVOCs) emitted by plants (e.g., isoprene, monoterpenes, sesquiterpenes) and anthropogenic VOCs, and there are estimated to be 10 000 to nearly 1 million different organic compounds in the atmosphere . VOCs are oxidized in the atmosphere by several different chemical species, such as hydroxyl, nitrate, chlorine radicals, and ozone . The oxidation of VOCs in the atmosphere leads to the production of SOAs. Presently, biogenic SOAs make up the majority of SOA loading (∼88 TgCyr-1) . The oxidation reactions are typically initiated at a double bond in the VOC, creating peroxy radicals . These peroxy radicals (RO2) can participate in further oxidation reactions, producing a range of products, such as organic nitrates (RONO2) and hydroperoxyl radicals (HO2) . Given this complex web of chemical interactions, computational models are crucial for understanding the climate effects of tropospheric chemistry.

While 3-D chemical transport models (CTMs), like GEOS-Chem, do exist and provide comprehensive organic chemistry, the chemistry by design is decoupled from the atmospheric dynamics: the transport model is usually run offline with prescribed wind fields, and chemical constituent output is then prescribed in an Earth system model (ESM) or general circulation model (GCM) . Extremely complex atmospheric oxidation chemistry models also exist, like the Generator for Explicit Chemistry and Kinetics of Organics in the Atmosphere (GECKO-A), and they provide detailed explicit chemistry of organic species and multigenerational chemistry . However, GECKO-A is typically run as a box model due to its computational expense.

In the past few years, however, several models have been developed that allow for comprehensive treatments of both tropospheric chemistry and other aspects of the Earth system while maintaining full coupling between the model components e.g.,. One such model is the Community Earth System Model version 2 with the Whole Atmosphere Community Climate Model version 6 as its atmospheric component (CESM2-WACCM6). Previous versions of this model included a more limited representation of organic chemistry compared to that in CESM2-WACCM6 . Climate simulations with more comprehensive chemistry typically prescribe sea surface temperatures (SSTs) e.g.,, which facilitates calculating standard climate change metrics like effective radiative forcing (ERF) e.g.,, but this approach limits the ability of the model to simulate climate responses to atmospheric forcings. CESM2-WACCM6, however, is capable of including both a dynamical ocean and comprehensive VOC chemistry in the troposphere. This model simulates interactive chemistry between organics and other atmospheric constituents, including aerosols (sulfate aerosol, salt, dust, etc.). These improvements to tropospheric chemistry also have an impact on gas-phase products such as O3 and NOx (NO + NO2).

These recently developed modeling capabilities present a new opportunity to investigate the climate effects of tropospheric VOC chemistry in a fully coupled model framework. In this study, we take advantage of this opportunity by comparing a control simulation of the standard CESM2-WACCM6 with a control simulation using a simplified CESM2-WACCM6 configuration in which detailed VOC chemistry is not included and emissions of a single bulk gas-phase SOA precursor are prescribed (referred to as a “SOAG scheme”). The difference between these two simulations provides an indication of the climate effect of tropospheric VOC chemistry. Past studies have compared the climates produced by CESM2 configurations with and without comprehensive chemistry in the troposphere . However, those CESM2 configurations also differed in their stratospheric chemistry, the number of model levels, the altitude of the model top, and the treatment of the quasi-biennial oscillation (QBO). This study is distinct in that it more precisely isolates the effect of tropospheric VOC chemistry without other confounding factors.

have compared configurations of CESM2 that also precisely isolate the effect of tropospheric VOCs, but their simulations used prescribed SSTs and sea ice, which, as mentioned before, limits the model's ability to produce a full climate response. Nonetheless, the findings of are very helpful for interpreting results of our simulations, as we will discuss below. Most of the simulations performed under the Aerosol Chemistry Model Intercomparison Project (AerChemMIP) also used prescribed SSTs . Among these prescribed SST simulations, there may be simulations that allow for assessment of the effects of VOC chemistry in other models (albeit without simulating the full climate response to VOC chemistry), and we plan to investigate those simulations further in the future. The AerChemMIP simulations that did not prescribe SSTs included experiments in which aerosols were prescribed and interactive. Comparison of such simulations, while also of great interest, would not isolate the effect of VOC chemistry, which is our focus here. Aerosols are prognostic in all of our simulations, but our model configurations differ in the chemical details with which aerosols are produced.

This study complements the extensive research focusing on the effects of stratospheric chemistry and particularly stratospheric ozone e.g.,. When the effects of tropospheric VOC chemistry are isolated, our simulations show significant impacts on climate, comparable in some regions to the temperature changes expected from anthropogenic greenhouse gas (GHG) increases over a century in future climate change scenarios. Even though only the tropospheric chemistry differs between our simulations, we will show that there are significant impacts over both the troposphere and stratosphere, and there are significant impacts on species that are not directly influenced by VOC chemistry, such as black carbon (BC) and primary organic matter (POM). In particular, we find that, compared to the SOAG scheme, the explicit VOC chemistry in CESM2-WACCM6 results in sulfate aerosol increases throughout the stratosphere and generally cooler temperatures in the extratropics. We will present evidence suggesting that these sulfate changes are primarily influenced by hydroxyl and SOA changes.

Despite the comprehensive chemistry in CESM2-WACCM6, there are (as detailed below) important limitations in this aspect as well as other aspects of the model (such as convective parameterization). So we cannot be certain that the climate effects highlighted in this study represent the full effects of tropospheric VOC chemistry in the real world. Rather, we consider this study to be a first step in disentangling the possible climate effects of VOC chemistry, which can help highlight areas for future model development and observational studies.

We begin by describing CESM2-WACCM6 and our experimental setup (Sect. ). Then we show the response to VOC chemistry, first focusing on the spatial structure of changes in key climate parameters (Sect. ), followed by investigation of the zonal mean response (Sect. ). The possible constituents responsible for the climate response are discussed in Sect. , along with the possible mechanisms responsible for changes in those constituents. Further details regarding cloud effects are examined in Sect. , and Sect.  provides a summary and concluding remarks.

What are the effects of tropospheric VOC chemistry on basic climate characteristics, such as surface temperature, clouds, precipitation, and winds? Figure  depicts the mean difference (WACtl - MACtl) in surface temperature, with red shading indicating warming and blue shading indicating cooling. Between New Zealand and South America, there is 0.5–1.5 K warming with 1–4 K cooling over the rest of the Southern Ocean. There is warming ranging from 0.5 to 2 K from Siberia across the Arctic Ocean through to the Barents and Greenland seas, with weaker warming over central Asia and central North America. Surface temperatures cool 0.5–1.2 K off the coast of Japan, with accompanying warming over the Gulf of Alaska. There is more localized strong warming (3–4 K) in southeastern Asia, the southwestern US, and the Great Lakes region. Localized pockets of cooling are interspersed with pockets of warming over central and southern Africa, eastern Australia, and central South America. The warming in particular regions is comparable to that generated by anthropogenic GHG increases over a century in future climate change scenarios e.g.,, which motivates further research to examine possible effects of VOC chemistry on anthropogenic climate change. While significant changes in surface temperature are seen over land, the localized mechanisms behind these changes are not clear and will be investigated in future studies. Overall, the surface temperature response shows no resemblance to the La Niña-like difference in the initial states of WACtl and MACtl (see discussion in Sect. ). This finding further establishes that the WACtl - MACtl responses are not being influenced by differences in their initial states.

The mean difference in global mean surface temperature (GMST) between WACtl and MACtl is -0.0640 K over the 211-year analysis period. This result indicates that, while VOC chemistry has a slight global cooling effect, there are much larger regional temperature effects that strongly cancel each other when averaging spatially. When imposing a climate change forcing, it is common to calculate the ERF to assess the strength of the forcing compared to forcings from other sources. A common method for calculating the ERF is the regression method described by . However, there is an initial warming of GMST in the first ∼ 20 years of WACtl relative to MACtl (for reasons that are unclear), even though the steady-state GMST response is negative, suggesting that the transient and equilibrium radiative forcings are opposite in sign. In such a situation, the regression method is not suitable for calculating the ERF, and simulations with fixed SSTs should instead be used e.g.,. Such simulations along with further investigation of the contrasting transient and equilibrium responses will be left for future work.

Taking a difference between our WACtl simulation and a 211-year segment of the MACtl-CMIP6 simulation (which, as detailed in Sect. , includes an interactive QBO) produces nearly identical results to our Fig.  (not shown), which suggests that the treatment of the QBO has a negligible effect on the surface temperature response. Given that the climates of MACtl and MACtl-CMIP6 are essentially the same, the 500 years of MACtl-CMIP6 output also allows for additional assessment of the noise associated with internal climate variability. We estimate this noise by averaging the surface temperature over a 211-year slice of MACtl-CMIP6 and then subtracting the surface temperature averaged over a non-overlapping 211-year slice of the same simulation. Figure  shows these time slice differences for four different choices of time range, obtained by shifting the beginning of each time slice by 20 years, with the results plotted on the same shading scale as in Fig. . The resulting differences for all of the time slice choices are negligible compared to the significant temperature differences shown in Fig. . In addition to our statistical significance tests, this analysis further demonstrates the robustness of our results.

A key influence on regional climate is the large-scale atmospheric circulation, which can be partially depicted through sea level pressure (SLP). Figure b shows the MACtl climatology of sea level pressure along with the approximate location of the midlatitude jets (in green), defined as the latitude where there is a local maximum in surface wind speed. In accordance with geostrophic balance, the midlatitude jets form on the boundaries between the polar lows and subtropical highs. Figure a shows the mean difference in SLP, with blue shading indicating decreases and red shading indicating increases. Over the South Pacific between New Zealand and South America, there is a weakening of the subtropical high accompanying an overall weakening of the polar low. These changes indicate a weakening of the meridional pressure gradient, which we expect to result in a weakening and equatorward shift of the midlatitude jet. (This shift will be confirmed further below when examining zonal mean zonal wind.) There is some strengthening of the subtropical high over the Indian Ocean, but this SLP change is smaller than the weakening of the polar low, and the meridional pressure gradient still decreases. Throughout the Northern Hemisphere (NH), the subtropical highs and polar lows weaken, also suggesting a weakening and equatorward shift of the midlatitude jet.

Regional precipitation changes can be due to changes in thermodynamics as well as dynamics. Figure  shows the mean difference of daily precipitation, with blue shading indicating increases and red shading indicating decreases. There are especially strong changes in precipitation around the Intertropical Convergence Zone (ITCZ): precipitation increases up to 1.5 mmd-1 over the northern and eastern portions of the Maritime Continent, with accompanying decreases (0.5–1 mmd-1) over the central and eastern equatorial Pacific. This band of decreased precipitation extends eastward to central Africa.

The precipitation changes over the equatorial Pacific are what would be expected with a strengthening of the Walker circulation e.g.,. Such a strengthening of the Walker circulation occurs under La Niña, along with cooling in the eastern equatorial Pacific e.g.,. However, Fig.  does not show significant cooling in the eastern equatorial Pacific, which suggests that other processes are responsible for the strengthened Walker circulation, a matter that requires further investigation.

Accompanying the precipitation decreases along the Equator, there is a band of increased precipitation just north of the Equator extending from the Atlantic eastward to Eritrea, indicating a northward shift of the ITCZ. Over the Americas, the equatorial precipitation decrease is accompanied by precipitation increases to the north and south, indicating a widening of the ITCZ. Over the Indian Ocean, precipitation increases over the Equator with accompanying decreases to the north and south, indicating a narrowing of the ITCZ. ITCZ shifts have been shown to relate to changes in the global energy transport e.g.,. Given the apparent regional variability in the ITCZ changes in our simulations, there are likely region-specific changes in energy transport at work see, which require further investigation.

There are precipitation decreases over most of the Southern Ocean (up to 0.7 mmd-1). These changes align with SLP increases over the Southern Ocean (Fig. a), with which we expect anomalous descent and decreased precipitation. Precipitation decreases over the Indian subcontinent, which corresponds to a weakening of the north–south pressure gradient in this region (Fig. a), indicating weaker flow from the Arabian Sea toward the Himalayas. Off the east coast of Japan, there are precipitation decreases of 0.1 to 0.4 mmd-1, which align with surface temperature cooling in that region (Fig. ), suggesting a primarily thermodynamic effect.

Over the subtropical South Pacific, roughly from Samoa to southern Chile, there is an arc of increased precipitation (up to 0.25 mmd-1) with an accompanying decrease to the south and west. There are qualitatively similar dipole changes in precipitation over the subtropics in the North Pacific and Atlantic. These subtropical precipitation changes do not align with significant temperature anomalies, indicating that they are likely driven by dynamical changes, such as changes in the Hadley circulation. Such changes cannot be related simply to SLP changes, and we will investigate the relevant dynamical changes further below.

Changes in clouds have a significant influence on the surface energy budget and hence temperature. Figure a shows the mean difference in total cloud fraction, and Fig. b shows the mean difference in net cloud radiative effect (net CRE), defined as the difference between the all-sky and clear-sky net radiation at the top of the atmosphere. Blue shading indicates decreases, while red shading indicates increases. Over most of the extratropics, cloud fraction changes and CRE changes are opposite in sign, as is expected in regions where low clouds predominate. In regions where there is large-scale subsidence, such as the eastern equatorial Pacific and subtropical dry zones, clouds are also confined to lower altitudes, and cloud fraction and CRE changes are opposite in sign.

In tropical regions of deep convection, however, such as the western equatorial Pacific and the equatorial Indian Ocean, there is not such a clear correspondence between cloud fraction and CRE changes, as there are competing effects of low and high cloud changes. As was found previously when examining precipitation changes (Fig. ), the cloud fraction changes over the equatorial Pacific show evidence of a strengthened Walker circulation, with cloud fraction increasing in the west and decreasing in the east. The dipole cloud fraction changes in the subtropics suggest a possible influence of changes in the Hadley circulation, which will be considered further below.

Over the Southern Ocean between New Zealand and South America, there are moderate (up to 3 %) decreases in the total cloud fraction with 2 to 7 Wm-2 increases in CRE. This CRE increase shows a spatial correspondence with warming over this region (Fig. ), suggesting that this warming is due to positive cloud feedbacks. The deceleration and equatorward shift of the midlatitude jet would be expected to reduce cloud cover in this region, and we will discuss this possibility further below.

Elsewhere over the Southern Ocean, there are moderate increases (up to 4 %) in the total cloud fraction with CRE decreases of up to 8 Wm-2. Intriguingly, SLP increases throughout this region (Fig. ), with which one might expect anomalous descent and reduced cloud cover. This result suggests that the cloud increases in this region are not dynamically driven and are instead a thermodynamic result of cooler temperatures resulting in a downward shift of the clouds. We will explore this possibility further below. Also, the surface temperature cooling in this region extends beyond the regions of CRE decrease. For example, there is cooling between Madagascar and Australia (Fig. ), despite a CRE increase in this region. This provides evidence that the cooling over the broader region encompassing the eastern Southern Ocean is not primarily driven by cloud feedbacks, although cloud feedbacks appear to amplify the cooling here. Furthermore, if atmospheric dynamics were driving the temperature changes here, one would expect either a dipole surface temperature anomaly or warming associated with anomalous adiabatic descent, neither of which is evident in Fig. . Overall, these results suggest that the cooling over this region is primarily due to non-cloud radiative effects rather than dynamical effects.

South of 60∘ S and from 150∘ E eastward to 60∘ W, there are decreases in CRE, and between 60∘ W eastward to 150∘ E there are increases in CRE. These changes show correspondence with changes in sea ice concentration (SIC; Fig. ), which changes in accordance with the surface temperature changes (Fig. ). That is, the CRE decreases (increases) here closely correspond to SIC decreases (increases). This makes sense because, if low cloud cover is fixed (which is approximately true around Antarctica) and SIC decreases (increases), then the surface albedo decreases (increases), and the radiative effect of low clouds gets stronger (weaker), implying a negative (positive) CRE change. The correspondence between CRE and SIC changes is less apparent around the date line, and consideration of other factors such as sea ice albedo and thickness changes might be required here, a topic for future work.

Over most of the NH extratropics, there are increases in cloud fraction (up to 3 %) with corresponding CRE decreases of up to 3.5 Wm-2. However, surface temperatures warm over most of this area (Fig. ), suggesting that clouds are not driving these temperature changes. The warming over the northern extratropical oceans shows some correspondence with SLP increases (Fig. ), suggesting that adiabatic warming is playing a role, and we would expect the warming of these bodies of water to generate warming of the surrounding land regions. In particular, the decreased SLP over Eurasia combined with the increased SLP over the pole would favor advection of warm anomalies from the Arctic Ocean toward Eurasia. We also cannot rule out the possibility that non-cloud radiative processes are contributing to the warming in the northern extratropics, and further investigation of the detailed causes is required.

The previous figures provided a global picture of the impacts of tropospheric VOC chemistry, and we can gain additional insight by examining the zonally averaged vertical structure of changes in key parameters. Figure  shows the mean difference of zonal mean atmospheric temperature, with red shading indicating increases and blue shading indicating decreases. There is 0.25–0.5 K cooling throughout the tropical stratosphere and Southern Hemisphere (SH) midlatitude troposphere. There is weaker cooling over most of the NH troposphere, which contrasts with surface temperature warming over most of the NH (Fig. ). Less apparent in Fig.  is that there is 0.05–0.31 K warming in the lowest three model levels north of 60∘ N. So at these latitudes, there is a transition from warming in the boundary layer to cooling in the free troposphere.

As noted previously, changes in large-scale dynamics play a key role in shaping the climate response to VOC chemistry. Figure a gives further detail of the dynamical changes by showing the mean difference of zonal wind, with red shading indicating increases and blue shading indicating decreases. In the SH, the stratospheric polar vortex weakens, and this negative zonal wind anomaly extends down into the troposphere on the poleward flank of the midlatitude jet. There is an accompanying acceleration on the equatorward flank of the SH midlatitude jet, implying an equatorward shift of the midlatitude jet, further confirming the shift suggested by the SLP changes examined previously (Fig. ). These changes follow what is expected from thermal wind balance: the zonal wind anomalies correspond to warming in the SH high latitudes which, in combination with midlatitude cooling, implies a reduction in the meridional temperature gradient. In the NH, there is a qualitatively similar weakening of the stratospheric polar vortex and zonal wind deceleration on the poleward flank of the midlatitude jet. However, the zonal wind response is quantitatively weaker in the NH than in the SH, and there is no clear indication of an equatorward shift of the NH midlatitude jet. This weaker zonal wind response in the NH compared to SH corresponds to the weaker tropospheric cooling and polar stratospheric warming in the NH.

Care is needed when comparing our zonal wind responses to those in earlier studies, since the atmospheric circulation response is highly sensitive to the regional structure of temperature changes. For example, idealized simulations have shown that a narrow band of warming around the Equator produces equatorward shifts of the midlatitude jets, whereas warming spread over all latitudes produces poleward shifts . In our simulations, the zonal wind responses in both hemispheres qualitatively resemble the circulation responses to high-latitude warming found in earlier studies e.g.,, and we will explore the mechanisms generating this warming further below. Moreover, these responses are qualitatively opposite to the poleward midlatitude jet shifts produced in simulations of global warming (with warming over all latitudes), including simulations in which SST is uniformly increased e.g.,. Based on such past work, we would expect that, aside from any regional temperature changes, widespread tropospheric cooling would also shift the midlatitude jets equatorward. Indeed, simulations of the response to increased natural aerosols have shown widespread tropospheric cooling and equatorward shifts of the midlatitude jets e.g.,. Such an effect might be contributing in the NH, where the tropospheric cooling is more spatially uniform than in the SH.

Interestingly, the zonal wind increase in the SH midlatitudes lies between the subtropical and midlatitude jets, which in combination with the weaker zonal wind acceleration in the deep tropics indicates that the SH subtropical jet is shifting poleward, opposite to the direction of the midlatitude jet shift. As the subtropical jet approximately corresponds to the edge of the Hadley circulation, this result suggests that the Hadley circulation is changing in ways that contrast with typical expectations based on the midlatitude circulation changes. This result motivates further examination of the mean meridional circulation (MMC), shown in Fig. b, with red shading indicating positive (clockwise) anomalies and blue shading indicating negative (counterclockwise) anomalies. The MMC is quantified using the meridional mass streamfunction, calculated by vertically integrating the zonal mean meridional wind Sect. 6.3.2. This figure confirms that there is slight expansion of the Hadley circulation in both hemispheres: there is a counterclockwise anomaly extending poleward of the SH HC edge and a clockwise anomaly extending poleward of the NH HC edge.

We have further confirmed these circulation changes using the metrics described in . Specifically, the HC poleward edges in both hemispheres were calculated by finding the zero crossing of the mass streamfunction at 500 hPa. This metric indicates a slight expansion of the HC in both hemispheres (0.0916∘ in the NH, 0.0483∘ in the SH), though these changes are not statistically significant. As noted previously when examining Fig. , the subtropical high strengthens over the Indian Ocean and weakens elsewhere, which suggests that the HC changes over the Indian Ocean might contrast with those at other longitudes. Indeed, when the longitudes corresponding to the Indian Ocean (20–150∘ E) are excluded from the mass streamfunction calculation, there is stronger, statistically significant HC expansion of 0.5990∘ in the SH (not shown). Furthermore, the midlatitude jet latitude (the local maximum of zonal wind at the lowest model level) shifts equatorward 0.3868∘ in the NH and 0.6231∘ in the SH, although the NH shift is not statistically significant. The SH jet shift is about half that produced over a century in climate change simulations e.g.,, motivating further research to investigate the possible effects of VOC chemistry on the large-scale circulation response to anthropogenic climate change.

Figure b also shows that the HC weakens in both hemispheres, although the weakening in the NH is much stronger than in the SH. In general, weakening of a particular overturning cell is indicated by an anomaly over the central portion of an overturning cell that is opposite in sign to the climatological streamfunction. In this case, HC weakening in the NH is indicated by a negative (counterclockwise) anomaly that opposes the positive (clockwise) HC climatology, and HC weakening in the SH is indicated by a positive (clockwise) anomaly that opposes the negative (counterclockwise) HC climatology. The counterclockwise anomaly over the NH HC crosses the Equator, indicating a northward shift of the rising branch of the HC, which aligns with the widespread precipitation decreases over the Equator and precipitation increases north of the Equator (Fig. ). This weakening and expansion of the HCs might help to explain some of the subtropical precipitation changes in Fig. . In particular, one would expect that, with HC expansion, there would be a precipitation decrease poleward of the original HC edge, and the HC weakening would lead to a precipitation increase in the subtropical dry zones, both of which are evident in the subtropical Atlantic and Pacific. In addition to these HC changes, Fig. b also shows weakening of the Ferrel cells in both hemispheres and weakening of the polar cell in the SH. As mentioned above, this weakening can be inferred from anomalies that are opposite in sign to the climatology.

To what extent are the temperature changes discussed above radiatively driven versus dynamically driven? Figure  shows the mean difference in heating rates due to various sources, with red shading indicating increases and blue shading indicating decreases. The total radiative heating (Fig. a) is calculated by summing the longwave heating (Fig. c) and shortwave heating (Fig. e). In a steady state, we expect the changes in total radiative heating to be balanced by the changes in dynamical plus latent heating (Fig. b) (here, dynamical heating was computed using the WACCM6 variable “DTCORE,” which includes both horizontal and vertical temperature advection). Comparing Fig. a and b, this steady-state balance is indeed satisfied. The climatological values of longwave heating are negative (except near the tropical tropopause), which means that positive changes (shaded red) typically indicate reductions in longwave cooling and negative changes (shaded blue) typically indicate increases in longwave cooling. Comparing Fig. a, c, and e, we see that most of the changes in total radiative heating are quantitatively accounted for by changes in longwave cooling.

In the stratosphere, the longwave cooling is mostly determined by local changes in blackbody emission, which follows the Stefan–Boltzmann law. Comparing Figs. c and in the stratosphere, we accordingly see that regions of increased (decreased) longwave cooling correspond to regions of increased (decreased) temperature. Near the surface, the longwave cooling is determined mostly by the difference between the downward longwave flux and the blackbody emission, with the latter dominating. Thus, over most of the lower troposphere, we also see reductions in longwave cooling corresponding to cooler temperature. There are parts of the extratropical mid-troposphere where, despite the reduced blackbody emission due to lower temperatures, the longwave cooling still increases. This change likely results from decreased longwave absorption due to the reduced upwelling longwave from the lower troposphere.

Comparing Fig. c and e, we see that most of the reduced longwave heating in the tropical stratosphere is due to decreased shortwave heating. Elsewhere in the stratosphere, as well as in the tropical upper troposphere and Arctic troposphere, the changes in longwave cooling are mostly explained by changes in dynamical heating (Fig. f). The anomalous dynamical heating is especially strong in the southern polar stratosphere (SPS), and it dominates over the increased shortwave heating. The increased shortwave heating here appears to be due to increased ozone concentration (Fig. ). In percentage terms, the largest ozone increases are in the troposphere, especially the ∼ 25 % increases in the tropical and SH upper troposphere. But in absolute terms, the largest increases (∼2.5×1011 molec. cm-3) are in the SPS. The reasons for these ozone changes in the stratosphere and troposphere will be investigated in a separate study. These results can be compared with earlier modeling studies which show that, when stratospheric ozone is increased on its own, the warming in the SPS is dominated by increased shortwave heating e.g., rather than dynamical heating. Altogether, these results suggest that the strong SPS warming in our simulations is dynamically rather than radiatively driven.

Lower in the troposphere, changes in latent heating (Fig. d) also have to be considered. Over most of the low and middle troposphere outside the Arctic, changes in latent heating are qualitatively opposite to changes in dynamical heating. This correspondence suggests that dynamical heating over most of the troposphere is due to vertical motions, as anomalous upward (downward) motion in the troposphere would be expected to produce anomalous adiabatic cooling (warming) as well as increased (reduced) condensational heating. The cancellation, however, is not perfect, and when latent heating and dynamical heating are summed together (Fig. b), it is clear that, over most of the troposphere, the changes in longwave cooling are mostly explained by changes in the sum of latent and dynamical heating. Hereafter, we refer to this combination of dynamical and latent heating as “dynamical–condensational” heating.

As noted above, there is warming near the surface in the polar regions, in contrast with the cooling over most of the extratropical troposphere. Over the Arctic, there is increased longwave cooling around 700 hPa, which we would expect to result in increased downwelling longwave at the surface. Figure  suggests that this increased longwave cooling aloft is due to increased dynamical heating, which would be expected with the anomalous descent resulting from the weakening polar low noted above (Fig. ). (Shortwave heating decreases in this region, so shortwave fluxes do not appear to be a dominant factor here.) As for the warming near Antarctica, longwave cooling decreases in this region despite the increase in blackbody emission, which suggests that reduced longwave cooling in the upper troposphere is dominating the longwave balance. However, such a change on its own would reduce surface temperature, which suggests that longwave changes do not explain the warming here. Instead, this warming is likely explained by shortwave heating, which does increase here. (This increased shortwave heating is more clearly visible in Fig. , which plots the shortwave heating on a linear pressure scale with an adjusted shading scale.) As we will show below, there is not a significant change in CRE here, so the shortwave heating changes near Antarctica are likely due to non-cloud shortwave forcers.

The results above show a strong role for dynamical–condensational changes in driving the longwave cooling changes and temperature response. However, given that the temperature response must ultimately be due to changes in VOC chemistry, there must ultimately be a change in radiative heating that kicks off the dynamical changes. The very close correspondence between the longwave and dynamical–condensational heating changes suggests that the longwave changes are likely dynamically driven. The complex structure of the longwave heating changes further argues for their dynamical origin: longwave radiative forcers are typically well-mixed and would be expected to produce more spatially uniform longwave cooling changes e.g.,see their Figs. 1 and 6. We have also examined changes in methane and carbon dioxide (Fig. ), and the changes are minuscule in percentage terms and of the wrong sign to explain the temperature responses in our simulations.

The above considerations suggest that the ultimate drivers of the response to VOC chemistry are shortwave forcers, which is expected since VOC chemistry primarily influences tropospheric aerosol formation. As noted above, earlier studies suggest that shortwave forcers are likely not driving the dynamical response in the extratropical stratosphere. However, in the tropical stratosphere, dynamical heating changes are small, and reduced shortwave heating is the dominant driver of the reduced longwave cooling. Thus, one possibility is that a change in a shortwave forcer in the tropical stratosphere drives the tropical stratospheric cooling, which then drives the dynamical response in the troposphere. Indeed, have shown that when cooling is imposed in the tropical stratosphere, it results in equatorward shifts of the midlatitude jets, in agreement with our results. However, such a forcing also produces HC contraction and weakening, which is opposite to the responses we obtain in our simulations. This contrast suggests that changes in shortwave heating in the troposphere are likely also contributing to the dynamical changes.

Such a possibility is surprising since the shortwave heating changes are much smaller than the dynamical–condensational changes over most of the atmosphere. Indeed, we previously used this fact to argue that shortwave heating changes were likely not driving the dynamical heating changes in the SPS. But given the stronger effects of dynamical motions, especially vertical motions, in the troposphere compared to the stratosphere, we cannot rule out the possibility that a change in shortwave heating in the troposphere might ultimately produce dynamical–condensational heating changes that are quantitatively larger than the shortwave heating change. The idealized simulations of provide support for such a possibility: in the upper troposphere, the warming due to the direct effect of an increase in the solar constant is about a factor of 4 smaller than the warming due to the changes in convective heating (compare their Figs. 3 and 8). Thus, it is worth taking a closer look at the shortwave heating changes in the troposphere.

Figure  shows decreased shortwave heating over most of the extratropical troposphere. Such reductions in shortwave heating could indeed generate cooling in the extratropical troposphere and result in dynamical changes (e.g., equatorward shifts of the jets, weakening of the meridional pressure gradient) that reinforce that cooling. Furthermore, there are more complex changes in radiative heating in the tropics that might explain some of the complexity in the circulation response (i.e., the opposite directions of the midlatitude and subtropical jet shifts).

What chemical species are responsible for these changes in shortwave heating? Addressing this question with certainty requires performing controlled experiments, which we plan to do in the future. For the current study, we have examined the patterns of changes in various constituents in order to document key changes and generate compelling hypotheses that can be tested with controlled experiments in the future. After examining numerous possible shortwave forcers, we found that changes in sulfate aerosol, BC, and POM are likely contributing to the widespread extratropical tropospheric cooling in our simulations.

VOC chemistry dramatically impacts total SOA concentration. Pure SOAs typically scatter incoming radiation, and increases in SOA burden represent a negative radiative forcing in future climate scenarios . Observational studies suggest that SOAs can also be absorbing aerosols (AAs) if internally mixed with BC , and such mixing effects are included in CESM2-WACCM6 . Indirect effects of SOA are also highly uncertain, with one study estimating a 0.02 ± 0.04 Wm-2 historical radiative forcing due to increased cirrus cloud formation when using an earlier version of CESM with a different aerosol scheme . As mentioned above, CESM2-WACCM6 includes both direct and indirect radiative effects of SOAs, but the indirect effect in this version of the model has, to our knowledge, not yet been quantified. Furthermore, as mentioned above, SOAs and other aerosols in CESM2-WACCM6 participate in tropospheric heterogeneous reactions that can affect other aspects of atmospheric composition.

Figure a shows the mean difference in SOA concentration, with blue shading indicating decreases and red shading indicating increases. In the tropical upper troposphere, there is a 260 % increase in SOA concentration. also noted increased SOA in this region when comparing fixed SST simulations of CESM2-WACCM6 with and without comprehensive VOC chemistry. This similarity suggests that the WACtl - MACtl changes in SOA do not ultimately arise from surface temperature changes. Rather, attributed this difference to the faster SOA formation in the SOAG scheme, which results in greater deposition in the lower troposphere and lower SOA concentration in the tropical upper troposphere compared to simulations with explicit VOC chemistry. There is also a ∼ 20 % SOA increase poleward of 50∘ N in the lower troposphere. Elsewhere in the troposphere, there are SOA reductions of up to 80 %, and nearly all of the stratospheric SOA is eliminated. We generally expect SOA increases as reactions are enabled between VOCs and SOA precursors. The reasons why SOAs decrease in some regions require more detailed investigation of changes in SOA precursors and their possible consumption by other chemical reactions, and we leave such investigation for future work. also noted a reduction in stratospheric SOAs due to fully interactive tropospheric chemistry. The SOA responses in our simulations do not offer additional direct insight into the shortwave heating responses, since as mentioned above, SOAs can have positive or negative effects on shortwave heating depending on the internal mixing. Nonetheless, Fig. a provides an indication of changes connected with numerous chemical pathways with potential climate effects that can be explored in the future.

The VBS mechanism for SOA formation would be expected to act as a sink for various radical species (OH, NO3, NO, NO2, and Cl). (To establish this effect with greater certainty, we would need to perform additional simulations that isolate the VBS mechanism, which is left for future work.) Reductions in these species can influence the HOx, NOx, and ClOx catalytic cycles, which in turn influence ozone formation . In particular, the reduction of tropospheric OH concentration due to the inclusion of isoprene chemistry has been well-documented in the literature e.g.,. To further illustrate just this one part of the complex web of interactions, Fig. b shows the change in OH concentration. This plot reveals an especially dramatic decrease in OH aligned with the strong increase in SOA in the tropical upper troposphere, which is expected since increased production of SOAs requires increased consumption of OH.

Increased emissions of VOCs (particularly isoprene) have been thought to increase the chemical loss of hydroxyl in the atmospheric boundary layer in low-NOx conditions . These initial hypotheses were contradicted by observational studies in tropical forests which observed elevated OH concentrations, suggesting a previously overlooked OH production term e.g.,. In order to reconcile these differences, modelers looked to incorporate mechanistic OH recycling in low-NOx conditions, which included isoprene epoxydiol formation, radical propagation in HO2–acylperoxy reactions, and H-shift isomerizations of isoprene–hydroxy–peroxy radicals . found that the H-shift isomerizations were an important mechanism for maintaining raised concentrations of OH in low-NOx conditions, with additional OH recycling mechanisms being identified in subsequent studies e.g.,. However, it has been shown that the instruments used in some of the previous observational studies may have produced a positive OH bias due to internally generated OH and oxidation of biogenic VOCs . Due to its short lifetime and high reactivity, the details of OH chemistry remain highly uncertain, and thus OH is represented in numerous ways across 3-D CTMs. More recently, found that the inclusion of isoprene oxidation in GEOS-Chem decreased the global mean concentration of OH by 11 %, with the largest decreases in the tropical continental boundary layer. The meridional structure of their OH changes qualitatively resembles that in our Fig. b, which suggests that inclusion of isoprene in WACtl leads to a reduction in OH compared to MACtl. also found that decreased OH increases the lifetime of CH4 and appears to impact the SO2–sulfate aerosol budget, as discussed further below.

BC has been classified as an AA, and it is treated as such in CESM2-WACCM6. BC can absorb incoming shortwave radiation and burn off clouds, a process known as the semi-direct effect. BC is thought to contribute to climate warming, though the exact magnitude of this effect is still uncertain . Figure a shows the mean difference in BC, with blue shading indicating decreases and red shading indicating increases. The MACtl climatology of BC (contours) peaks in the tropics and the NH midlatitudes, indicative of the rapid industrialization already occurring in Europe and North America during 1850. Figure  shows decreased BC throughout most of the atmosphere, with the largest decreases (20 %–60 %) occurring in the NH extratropical troposphere. We would expect such a decrease to contribute to the reductions in shortwave heating shown in Fig. , especially in the NH troposphere.

POM, in contrast to BC, is not strictly an AA. This ambiguity is due to the definition of POM. POM is considered aerosol that is directly emitted from diesel and gasoline exhaust as well as biomass burning (both anthropogenic and natural). POM can act as nuclei for SOA, or POM can become internally mixed with BC. POM can also be internally mixed with sulfate aerosol, which typically scatters more than it absorbs. So depending on the aging that occurs, POM can act more like an AA or more like a scattering aerosol . In our simulations, however, the spatial distribution of POM in MACtl (Fig. b) closely corresponds to that of BC (Fig. a), and the WACtl - MACtl reductions in POM (30 %–65 %) also closely correspond to those of BC. This correspondence suggests that, in our simulations, POM is internally mixed with BC and acts primarily as an AA. Thus, we can expect that the reduction in POM will also contribute to reductions in shortwave heating, especially in the NH troposphere.

also noted similar decreases in BC and POM when comparing fixed SST CESM2-WACCM6 simulations with and without explicit VOC chemistry. (The similarity of our results to those of , is especially clear when comparing absolute changes in BC and POM; not shown.) This similarity suggests that these changes are ultimately not due to surface temperature changes. Rather, link these changes to overall increased SOA formation, which results in increased internal mixing of POM and BC with SOA, increased aging of POM and BC into the hydrophilic accumulation mode, and greater wet deposition of BC and POM in simulations with VOC chemistry compared to simulations with the SOAG scheme. Thus, even though BC and POM are not directly influenced by VOC chemistry, the effect of VOC chemistry on SOA production has indirect effects on BC and POM levels. show, however, that MAM4 overestimates BC mixing with other aerosol types, so it is possible that the BC and POM responses in Fig.  are exaggerated, and additional work is needed to assess the indirect effects of VOC chemistry on BC and POM.

Increases in sulfate aerosol (primarily from volcanic eruptions, with smaller contributions from biomass burning and fossil fuel emissions) can lead to surface cooling, while decreases in sulfate aerosol can lead to surface warming . There are also biogenic contributions to sulfate aerosol, primarily dimethyl sulfide (DMS) from marine sources . Figure  shows the mean difference in sulfate aerosol, with red shading indicating increases and blue shading indicating decreases. Sulfate increases (up to 15 %) in the SH extratropical upper and lower troposphere. Elsewhere in the troposphere, sulfate decreases with an especially pronounced decrease of approximately 35 % in the tropical upper troposphere. This reduction somewhat aligns with the pronounced increase in SOA shown previously (Fig. a), suggesting that some of the sulfate aerosol reduction is due to internal mixing with SOA. This influence of internal mixing with SOA has also been noted by , and this also influences the BC and POM changes as discussed above. The tropical sulfate changes appear to have an even closer alignment with OH changes (Fig. b). Note especially the OH increase in the tropical tropopause and the OH decrease in the tropical upper troposphere that align with qualitatively similar changes in sulfate. This correspondence suggests that OH changes are the dominant driver of tropical tropospheric sulfate changes, and we consider this possibility further below.

Sulfate aerosol is essentially unchanged in the tropical and NH lower stratosphere. Elsewhere in the stratosphere, there are sulfate increases ranging from about 3 %–8 %. These increases approximately align with reductions in SOA (Fig. a), suggesting that reduced internal mixing with SOA is contributing here. However, other contributions, such as changes in the stratospheric circulation, need to be explored further, especially in the tropical and NH lower stratosphere, where sulfate is essentially unchanged despite the decreased SOA. We would expect the sulfate increases in the extratropical stratosphere to further contribute to cooling in the extratropical troposphere. Furthermore, the sulfate reductions in the tropical troposphere would, on their own, lead to warming in the tropical troposphere. Thus, it is possible that, in the tropical troposphere, warming due to decreased sulfate aerosols is opposing cooling due to changes in other constituents (e.g., BC, POM), resulting in nearly unchanged temperatures in the tropical troposphere (Fig. ). As mentioned above, verifying such a possibility will require additional controlled experiments, which we plan to perform in the future.

Figure a shows the zonal mean percent change in the vertical column density (VCD) of sulfate aerosol burden. These changes reiterate key features apparent in Fig. , showing that vertically integrated sulfate decreases in the tropics and increases in the extratropics. We have performed additional analyses to assess the roles of OH changes and wet deposition. The primary chemical sink of SO2 is due to a reaction with OH, a process that eventually produces sulfate aerosol . Sulfate can be removed from the atmosphere via wet and dry deposition , and its abundance is also influenced by internal mixing with SOA as discussed above. Therefore, if there were no changes in SO2, SOAs, or wet deposition, a decrease in OH would result in a decrease in sulfate.

Figure b shows that the VCD of SO2 increases from approximately 60∘ S to 40∘ N, whereas Fig. d shows that the VCD of OH decreases over all latitudes, with the strongest OH decreases in the tropics. This suggests (as was apparent from our earlier comparison of Figs. b and ) that the tropical sulfate aerosol decreases are primarily due to decreases in OH. These decreases appear to be partially diminished by a precipitation decrease south of the Equator (implying less wet deposition) and enhanced by a precipitation increase north of the Equator (implying increased wet deposition; Fig. a, c). This result indicates that wet deposition changes are also contributing to sulfate aerosol changes, but these changes appear to be secondary to the chemical influence of OH changes. As discussed above, tropical SOA increases (Fig. a) might also be contributing to the tropical sulfate decreases. Furthermore, the most plausible explanation for increased sulfate aerosol outside the tropics is the decreased extratropical SOA, as SO2 and OH changes are weak in the extratropics. The sulfate aerosol increase in the SH extratropics appears to be further enhanced by a decrease in precipitation over approximately 30–70∘ S.

In addition to the changes in shortwave forcers mentioned in the previous section, changes in clouds are also expected to contribute to changes in shortwave heating. In Sect. , we showed evidence that cloud feedbacks were influencing surface temperature changes over widespread regions. Here, we explore these changes further from the zonal mean perspective to assess possible contributions to zonal mean temperature changes.

Figure  shows the zonal mean differences in net CRE (panel a) along with the zonal mean changes in cloud fraction (panel b). (In panel b, red shading indicates increases in cloud fraction and blue shading indicates decreases.) There is an overall downward shift in clouds over most latitudes. Such a downward shift is expected, as cooler temperatures would result in condensation occurring at lower altitudes. Such a downward shift is qualitatively opposite to the upward shift seen in global warming simulations e.g.,. The extent to which changes in cloud condensation nuclei (CCN) might also be contributing to these cloud changes needs to be further investigated. The changes around 60∘ S are more vertically aligned, suggesting that dynamical changes, such as the equatorward shift of the midlatitude jet, are playing a role. Indeed, earlier studies have shown the strong connection between clouds and the response of the SH midlatitude jet to climate forcing e.g.,. In the SH midlatitudes and NH high latitudes, the downward shifts of the clouds result in CRE decreases as expected. As the clouds shift equatorward with the SH midlatitude jet, there is a sharp increase in CRE around 60∘ S, on the poleward flank of the midlatitude jet. Around this latitude, there is also a reduction in sea salt aerosol associated with the reduced near-surface wind speeds that appears to drive a reduction in cloud liquid (not shown) and amplify the cloud fraction decrease as well as CRE increase. Near the Equator, the downward shift in upper-level clouds appears to be counteracted by a reduction in mid-troposphere clouds, resulting in a CRE increase here.

Comparing the cloud fraction changes with the shortwave heating changes (Fig. ), we see a number of locations where the changes correspond to each other, such as in the tropical middle to upper troposphere and around 50∘ S (in the vicinity of the SH midlatitude jet). This correspondence suggests that cloud changes are driving some of the shortwave heating changes. The apparently cloud-driven shortwave heating increase spread over the tropical upper troposphere is of particular interest because idealized studies have shown that such heating on its own produces HC weakening and expansion e.g.,. Thus, the complexity in the circulation response due to VOC chemistry may be due to the competing effects of extratropical shortwave forcers and tropical cloud feedbacks. Widespread extratropical cooling (possibly due to changes in sulfate aerosols, BC, and POM) would be expected to shift the midlatitude jets equatorward, whereas tropical increases in shortwave heating (likely due to tropical cloud changes) would result in HC weakening and expansion.

This study examines the climate effects of tropospheric VOC chemistry by taking the difference between two simulations of CESM2-WACCM6: (1) WACtl, a case with explicit VOC chemistry, and (2) MACtl, a case with a simplified SOAG scheme and no explicit VOC chemistry. Changes (WACtl - MACtl) in surface temperature range between -4 and 4 K. This result motivates further research to investigate the possible influence of VOC chemistry on anthropogenic climate change. While we explored the mechanisms of the large-scale temperature changes, our simulations also show significant localized temperature changes over land. The reasons for these localized changes are unclear, and this is an important topic for follow-up study. Such an investigation would require analysis of regional changes in shortwave, longwave, and dynamical heating, going beyond our zonally averaged analysis of these quantities.

In the zonal average, there is widespread tropospheric cooling in the extratropics. This cooling results in an equatorward shift of the SH midlatitude jet that is in turn associated with shifts of precipitation patterns and cloud cover. This cloud shift results in a positive cloud feedback in the SH high latitudes that offsets cooling there. Elsewhere in the extratropics, the overall cooler temperatures (possibly in combination with CCN changes) produce a downward shift of clouds that produces additional cooling in the midlatitudes. There is also weakening and expansion of the HCs that appears to be due to a cloud-driven shortwave heating increase in the tropical upper troposphere. This HC weakening drives additional precipitation changes in the tropics and subtropics.

We presented evidence that changes in longwave forcers were likely not the root cause of these climate responses. Instead, there is evidence that shortwave forcers, such as decreased BC, decreased POM, and increased stratospheric sulfate aerosols, are contributing to the extratropical cooling. In the tropics, there is evidence that reductions in sulfate aerosols offset the cooling due to other shortwave forcers, resulting in nearly unchanged tropical temperatures. Confirming the roles of these shortwave forcers will require performing additional controlled experiments, which we plan to do in the future.

As found in , slower SOA formation allows for increases in SOA that are especially dramatic in the tropical upper troposphere. This overall SOA increase results in increased internal mixing of sulfate aerosol, BC, and POM with SOA, which strongly influences the changes in these constituents. Furthermore, SOA decreases throughout the stratosphere, which appears to drive the sulfate increase over most of the stratosphere. Thus, the effects of VOC chemistry extend to aerosols that do not directly participate in VOC chemistry.

Our simulations also produce significant changes in tropospheric and stratospheric ozone (Fig. ). We presented evidence suggesting that these ozone changes likely do not play a strong role in the large-scale climate response, as the strong warming in the SPS is mostly driven by dynamical heating. Nonetheless, we plan to perform additional controlled model experiments with prescribed ozone to confirm that ozone changes are playing a minor role. Furthermore, the mechanisms driving these ozone changes are interesting in their own right and deserving of further study.

There are important limits to the representation of chemistry in CESM2-WACCM6, such as the exclusion of HONO. Furthermore, for simulations that include recent anthropogenic emissions (which our simulations exclude), there are additional species absent from CESM2-WACCM6, such as methyl nitrate and peroxypropionyl nitrate (PPN), that might be important to consider. Thus, additional work is needed to improve the representation of chemistry in CESM2 and develop confidence that the responses we obtain capture the effects of VOC chemistry in the real world. As such improvements are made, maintaining a hierarchy of models with both coupled and uncoupled chemistry will enable continued efforts to disentangle the numerous possible climate effects of tropospheric chemistry. In addition to CESM2 improvements, we look forward to additional analysis of existing simulations, additional simulations with other models, and additional observational work that can help to assess the generality of our findings.