Significance A mass extinction occurred at the Cretaceous−Paleogene boundary coincident with the impact of a 10-km asteroid in the Yucatán peninsula. A worldwide layer of soot found at the boundary is consistent with global fires. Using a modern climate model, we explore the effects of this soot and find that it causes near-total darkness that shuts down photosynthesis, produces severe cooling at the surface and in the oceans, and leads to moistening and warming of the stratosphere that drives extreme ozone destruction. These conditions last for several years, would have caused a collapse of the global food chain, and would have contributed to the extinction of species that survived the immediate effects of the asteroid impact.

Abstract Climate simulations that consider injection into the atmosphere of 15,000 Tg of soot, the amount estimated to be present at the Cretaceous−Paleogene boundary, produce what might have been one of the largest episodes of transient climate change in Earth history. The observed soot is believed to originate from global wildfires ignited after the impact of a 10-km-diameter asteroid on the Yucatán Peninsula 66 million y ago. Following injection into the atmosphere, the soot is heated by sunlight and lofted to great heights, resulting in a worldwide soot aerosol layer that lasts several years. As a result, little or no sunlight reaches the surface for over a year, such that photosynthesis is impossible and continents and oceans cool by as much as 28 °C and 11 °C, respectively. The absorption of light by the soot heats the upper atmosphere by hundreds of degrees. These high temperatures, together with a massive injection of water, which is a source of odd-hydrogen radicals, destroy the stratospheric ozone layer, such that Earth’s surface receives high doses of UV radiation for about a year once the soot clears, five years after the impact. Temperatures remain above freezing in the oceans, coastal areas, and parts of the Tropics, but photosynthesis is severely inhibited for the first 1 y to 2 y, and freezing temperatures persist at middle latitudes for 3 y to 4 y. Refugia from these effects would have been very limited. The transient climate perturbation ends abruptly as the stratosphere cools and becomes supersaturated, causing rapid dehydration that removes all remaining soot via wet deposition.

The Cretaceous−Paleogene (K−Pg) boundary coincides with an asteroid impact and marks one of the five great extinction events since the Cambrian explosion of life forms 541 Ma. The millimeter-thick portion of the boundary layer far from the asteroid impact site at Chicxulub, in the Yucatán Peninsula, contains iridium, which was used to identify the asteroid impact at the time of the mass extinction event 66 Ma (1⇓–3). According to Wolbach et al. (4), it also contains as much as 56,000 Tg of elemental carbon, of which 15,000 Tg is in the form of fine soot nanoclusters, and the remaining 41,000 Tg is made up of coarser soot particles. Earlier estimates by the same authors (5), based on a smaller number of samples, yield even larger numbers: 70,000 Tg of soot, of which 35,000 Tg is fine soot. Although many details of the extinction event and the origins of various materials in the K−Pg layer are poorly understood, the presence of soot is incontrovertible. The soot is collocated with the iridium, and therefore must have been injected during the time required for the iridium to be removed from the atmosphere and reach the ground; it could not have come from forest fires decades or centuries after the impact (4). Although some argue that the soot originated from burning hydrocarbons at the impact site (6), recent studies indicate that the hydrocarbon source is quantitatively insufficient to explain the soot layer (7). The mass of soot is so great for the 70,000 Tg estimate that most of the aboveground biomass, and likely much of the biomass in the near-surface soil, must have burned immediately following the impact and produced fine soot with high efficiency (4, 8).

In this study, we present simulations of the short-term climate effects of massive injections of soot into the atmosphere following the impact of a 10-km-diameter asteroid. We assume that the soot originated from global or near-global fires (8). The short-term climate effects of the soot would augment and probably dominate those of other materials injected by the impact, which are not considered here except for water vapor. Given the range of estimates for the fine soot produced by the impact (4, 5), we consider soot injections of 15,000 Tg and 35,000 Tg. Substantially smaller estimates have been proposed (9), so we also simulate a much smaller soot injection, 750 Tg, to contrast the climate effects of large and small soot injections.

Materials and Methods We use the Community Earth System Model (CESM) (10), a fully coupled climate model that includes atmosphere, ocean, land, and sea−ice components. We use the Whole Atmosphere Community Climate Model, version 4, (WACCM) as the atmospheric component (11). WACCM is a “high-top” chemistry−climate model, with an upper boundary located near 140-km geometric altitude; it has horizontal resolution of 1.9° × 2.5° (latitude × longitude), and variable vertical resolution of 1.25 km from the boundary layer to near 1 hPa, 2.5 km in the mesosphere, and 3.5 km in the lower thermosphere, above about 0.01 hPa. WACCM is used as the atmospheric model to be able to simulate the physical and chemical consequences of injection and lofting of impact materials to great heights in the atmosphere. The upper range of the estimated soot burden produced by the asteroid impact is 70,000 Tg (5). To represent the evolution of such a massive injection accurately, we have coupled WACCM with the Community Aerosol and Radiation Model for Atmospheres (CARMA) (12). CARMA is a sectional aerosol parameterization that resolves the aerosol size distribution. CARMA aerosols are advected by WACCM, are subject to wet and dry deposition, affect the surface albedo, and are included in the WACCM radiative transfer calculation. The soot is treated as a fractal aggregate for both microphysics and radiative transfer (13), and coagulation of soot particles is considered. The fractal particles have a monomer size of 30 nm, a fractal dimension varying between 1.5 and 3.0, and a packing coefficient of 1 (13). The largest burdens of soot aerosol considered here cause enormous temperature changes in the stratosphere and mesosphere, which required changes to WACCM to improve the numerical stability of the model. These changes and additional details about the model configuration are described in Supporting Information. We carried out seven simulations for this study, a 20-y control simulation and six 15-y perturbation experiments, described below and summarized in Table S1. We also carried out a few additional short simulations with output at high temporal resolution to assess the impact of soot injections between 750 Tg and 35,000 Tg on solar flux at the surface. Data from the simulations will be made available on request. All simulations use modern continental positions and atmospheric composition. Initial conditions for the calculations are discussed by Toon et al. (8). Soot is assumed to be produced by global fires ignited as debris from the impact falls through the atmosphere at high velocities and heats up to very high temperatures. We assume that fine soot is lofted to the upper troposphere in pyrocumuli (14), and we thus place it in a Gaussian distribution centered on the local tropopause. The remaining, coarse soot particles are placed in a half-Gaussian at the surface. Both fine and coarse soot are injected over 24 h. The coarse soot is removed rapidly (Fig. S1) and plays a negligible role in forcing climate change; therefore, in the remainder of the paper, we refer to the various simulations by the amount of fine soot injected. See Soot Emission and Removal and Toon et al. (8) for a detailed discussion of the soot emissions. Table S1. Simulations carried out for this study Fig. S1. Evolution of soot burden (black) and cumulative wet (blue) and dry (red) deposition for the simulation with 15,000 Tg of fine soot and 41,000 Tg of coarse soot, including injections of H 2 O, CO 2 , and heat from fires for (A) 1 wk using hourly data and (B) 1 y using monthly average data. The soot is emitted over the first 12 h to 36 h and is followed by rapid wet and dry removal. After 1 y, the remaining soot burden is 4% of the total 56,000-Tg emission, with 57% removed by wet deposition and 38% removed by dry deposition. There are many other materials that plausibly might have been injected along with the soot (8). The K−Pg layer is dominated by spherules about 200 µm in diameter (15). These particles likely ignited the global wildfires, but they could have remained in the atmosphere only a few days and would not have impacted climate directly. In addition, clastics were clearly produced in the impact and extend over much of North America. However, the submicron fraction that could have been part of the global aerosol layer is subject to debate, difficult to determine from theory, and not detectable in the K−Pg global layer because of chemical weathering of the clastics. Vaporized impactor and target material would not only have condensed to form the large spherules but may also have left behind a large mass of rock vapor (16). The fate of the vapor is unknown; it may have condensed on the spherules, or it may have entered the atmosphere and recondensed there as nanometer-sized particles. Another impact material that might have been present but cannot be documented in the K−Pg boundary layer is sulfur originating from the asteroid or the target rock at the impact site. A number of authors (8, 17⇓⇓–20) have suggested that sulfur injections may have modified the climate after the impact. It is known that large volcanic eruptions that inject sulfur into the stratosphere can affect climate by reducing the solar flux that reaches the troposphere. It is not clear how much of this sulfur may have reacted on the spherules or rock vapor and then been quickly removed. The total amount of sulfur injected by the impact is estimated to be about 100 Gt (100,000 Tg) (17), which is larger than the mass of soot observed in the K−Pg layer. Even so, previous investigators (18, 19) found that the sulfate alone could not reduce light levels below 1%, because sulfate is nearly transparent at visible wavelengths. However, a recent study (20) using a climate model with the sulfate aerosol radiative forcing estimated by Pierazzo et al. (19) shows that reduction of the solar flux to a few percent of normal values is sufficient to cause severe global cooling. Thus, both soot and sulfate aerosols are sufficient to produce large, transient decreases in global temperature, but large injections of soot will also cause near-total darkness at the surface for a protracted period. Because of the uncertainties associated with the presence and possible impacts of materials other than soot, they are not considered in our simulations. However, we do consider the effects of injecting into the atmosphere, together with the soot, a large amount of water vapor produced by vaporized and splashed seawater at the impact site, and as a combustion product from the global fires. Toon et al. (8) estimated that 7.5 × 106 Tg of water was produced, with 1.5 × 106 Tg coming from combustion, and we use these estimates. Consideration of the effects of water vapor is important because such a massive injection would have produced supersaturated conditions above the tropopause. Subsequent rainout could then remove a possibly important fraction of the soot from the upper atmosphere. To put our results into context, we also carried out a simulation using a considerably smaller amount of fine soot, as done by Kaiho et al. (9), who assumed soot was produced from carbon present at the impact site, but estimated a much smaller soot input than Wolbach et al. (4, 5). In addition to a much smaller soot injection, Kaiho et al. did not include coagulation in their simulations, so their particles did not grow in time, and the size did not change with the mass injected. In their standard 1,500-Tg case they used an initial soot particle size mode of 11.8 nm, which is much smaller than smoke in the present-day atmosphere. Toon et al. (8) recommended an initial soot particle size mode of 110 nm, which is based on Wolbach et al.’s (21) analysis of the particle size in the K−Pg layer, and is also very similar to observations of modern forest fire smoke. The optical properties of 11.8-nm particles are much different from those of more realistic smoke particles. In all of our simulations, we inject the fine soot near the tropopause, with an initial size of 110 nm. We note, finally, that we have not included the effects of CO 2 release from the impact site, nor the CO 2 and heat of combustion from the burning of biomass in most of our calculations. The omission of CO 2 was dictated by technical considerations, as the parameterization of nonlinear thermodynamic equilibrium infrared transfer in our model was unstable for the very large mixing ratios of CO 2 produced following the impact. We return to this point in Discussion, where we show that neither a massive injection of CO 2 nor the heat from global fires affects significantly the short-term response to the asteroid impact.

Discussion According to Wolbach et al. (4), about 15,000 Tg of fine soot and 41,000 Tg of coarse soot is present in the millimeter-thick, global K−Pg boundary layer, which also contains iridium that identifies the layer as the result of an asteroid impact (1). The soot is believed to have originated from global fires (4, 5, 8). These fires would have been an efficient extinction mechanism for large land animals (7, 34). Placing 15,000 Tg of fine soot into our global climate model shows that 95% of the soot would be removed from the atmosphere in a year, defining the timescale that is represented by the layer in land deposits. On the other hand, it might take decades for the small soot particles to fall to the bottom of the oceans, assuming no zooplankton were present to consume it and excrete it in large fecal pellets. Sunlight absorption by soot reduces surface shortwave irradiance to levels lower than found today at the base of the ocean euphotic zone for a year or more, which would trigger a collapse of the ocean food chain and extinctions of marine organisms that depended on the photosynthetic productivity of the euphotic zone (31⇓–33). There are no mechanisms other than impacts that have been suggested to produce such low light levels in post-Cambrian Earth history. There are no refugia from the low light levels. However, even for a 15,000-Tg soot injection, light levels would rise above 1% of normal (the level below which photosynthesis is severely inhibited) in the Tropics after 2 y following the impact, and a year earlier in polar latitudes. The lack of sunlight leads to dramatic cooling of the planet, by over 15 °C on a global average, 11 °C over the ocean, and 28 °C over land. Global average temperature cooling of about 5 °C occurred during the last ice age relative to the warmest part of the Holocene (50). Therefore, the cooling following the K−Pg impact was larger than that in an ice age and much more sudden, but of much shorter duration (years vs. tens of thousands of years). Sudden ocean cooling is consistent with the TEX 86 paleo-sea surface temperature proxy record from Vellekoop et al. (51), which shows cooling of 7 °C over a period of months to decades postimpact. There were likely refugia from freezing temperatures on land in the Tropics, and along coastlines following the K−Pg impact. Temperatures in the ocean euphotic zone decline 10 °C on a global average in our simulations, but temperatures below 500 m depth are not affected, such that the deep ocean would have been a refuge from temperature changes. It should be noted that global cooling such as obtained for our 15,000-Tg soot case can also occur under different scenarios. For example, Brugger et al. (20) have calculated cooling of similar magnitude by assuming that the asteroid impact injected a very large amount of sulfur (100 Gton) into the stratosphere, which then formed sulfate aerosols and scattered sunlight. This amount of sulfur produces reductions in sunlight at the surface comparable to those obtained from a relatively small injection of soot, 750 Tg, because soot is a much more efficient absorber of sunlight. Interestingly, our calculations for a 750-Tg soot injection produce cooling comparable to our 15,000- and 35,000-Tg cases. In general, any mechanism that can reduce sunlight at the surface to a few percent of normal values for a protracted period is sufficient to induce severe cooling (tens of degrees Celsius) lasting as long as the sunlight-blocking material remains in the stratosphere. However, a large injection of sulfur cannot produce the near-total darkness, lasting for almost 2 y, that follows large injections of soot. Thus, a massive injection of soot into the stratosphere adds the effects of darkness on the food chain to the stresses of global cooling and decreased precipitation. Interestingly, a soot injection of only 5,000 Tg, 3 times smaller than the 15,000-Tg estimate of Wolbach et al. (4), still reduces light levels below 1% for 1 y. This result implies that extensive but less-than-global fires, such as suggested in some models for the distribution of impact debris (52, 53), would also suppress primary productivity for a prolonged period. While the surface and lower atmosphere cool due to screening of sunlight by the airborne soot, the tropopause warms by over 50 °C and the upper atmosphere by as much as 200 °C for a 15,000-Tg soot injection. The warm tropopause temperatures eliminate the tropical cold trap and allow water vapor mixing ratios to increase to well over 1,000 ppmv in the stratosphere. High stratospheric temperatures accelerate the destruction of ozone via the O + O 3 reaction, and large water vapor mixing ratios are a source of HO X radicals, which are efficient catalysts of ozone destruction. As a consequence of the enormous increases in temperature and water vapor following the impact, the ozone layer is partially removed for 7 y, with ozone column amounts dropping as low as 20% of normal on a global average. However, absorption of sunlight by the soot protects the surface from UV light, except for a period of about 2 y during the sixth to eighth years after the impact. During this period, UV light is able to reach the surface at highly elevated and harmful levels. High UV exposure at the surface ends when water, and therefore HO X species, is removed as the stratosphere cools and becomes supersaturated, resulting in an abrupt dehydration event during year 7 after the impact. Longer-term loss of ozone could have occurred via injection into the stratosphere of halogen species (Cl, Br) from splashed seawater, as shown by Pierazzo et al. (46), but these were not included in the present calculations. Toward the end of this investigation, we devised a solution to the numerical instability problem that occurred when simulating massive CO 2 increases and repeated the 15,000-Tg case adding an injection of 2.46 × 106 Tg of CO 2 (8), about 0.78 of the present-day atmospheric mass of this gas. Global average temperatures in the simulations with and without additional CO 2 diverge after the impact-generated soot is removed. After 15 y, the additional CO 2 produces a 1 °C increase in global surface temperature and a 5% increase in the global ozone column. Thus, CO 2 affects the longer-term evolution of the climate, but not its short-term response, consistent with the results of Brugger et al. (20) and the K−Pg temperature reconstruction from Vellekoop et al. (51). We also repeated the 15,000-Tg simulation including 4.6 × 1022 J of heat from combustion in global fires (Fig. S8 for details). During the fires, the global average surface temperature increases by 10 °C (Fig. S8A); however, after 3 y, the difference relative to the simulation that omits heat input from the fires is only about 1 °C (Fig. S8B). Fig. S8. Evolution of global average surface temperature from a simulation with 15,000 Tg of soot and heat from combustion by global fires (red), and a similar simulation without the added heat (black) using (A) hourly data for 1 wk and (B) monthly average data for 3 y. The heat from fires is estimated to be 4.6 × 1022 J based upon a land biomass of 2 g⋅cm−2 and the energy from combustion for air-dried wood, 1.5 × 104 J⋅g−1. Heat from the fires is emitted into the lowest three model layers (surface to ∼900 hPa), and the fires are active for the first 12 h to 36 h of the simulation. During this time, there is a sharp rise of 10 °C in surface temperature; however, the temperature difference quickly drops to ∼2 °C after a week. There is an increase in surface temperature of up to 3 °C during the first few months as the added heat is transferred from the atmosphere to the surface, but this is gradually reduced down to about 1 °C by the end of 3 y, as the loss of downwelling solar radiation dominates the surface energy budget, and some of the added heat radiates to space. The maximum cooling of 15 °C in this simulation occurs during the third year (Fig. 4), so the heating from combustion causes only a 7% reduction of this peak cooling. The sharp initial warming would be reduced if the fires were assumed to occur over a longer period. The transient consequences of a large asteroid impact followed by global fires, which include suppression of primary productivity during a protracted period of darkness, severe and widespread cooling at the surface, and high doses of UV radiation, appear to be enough to account for nearly instantaneous and widespread species extinction at the K−Pg boundary. Further work should consider Late Cretaceous geography and climate (54), halogen injections from seawater, and more-complex aerosol microphysics, including organic carbon, oxidation, and sulfate coatings.

SI Text Here, we provide additional background on how the soot emissions used in this study were determined, describe some of the characteristics of the aerosol microphysics, give details about how WACCM was modified to remain stable under the very large perturbations experienced in these simulations, and indicate how the calculations of the surface UVI were carried out.

Soot Emission and Removal We use the estimates of aerosol and tracer injections from the K−Pg impact described by Toon et al. (8). As regards the soot, we are guided by Wolbach et al. (4, 5), who measured the elemental carbon found in samples of the K−Pg boundary clay. Samples were first etched to remove kerogen (organic carbon), and the remaining insoluble carbon was then separated into samples, including small, “aciniform” soot clusters plus other larger charcoal or carbon particles. We use the terms “fine soot” and “coarse soot” to refer to these two categories. Using samples from five sites, Wolbach et al. (5) estimated that 70,000 Tg of elemental carbon is present in the K−Pg layer, of which 50% (35,000 Tg) is fine soot. With the addition of six more sites, Wolbach et al. (4) estimated total elemental carbon to be 56,000 Tg, of which 26.6% (15,000 Tg) is fine soot. They measured the size distribution of the fine soot particles but did not determine that of the coarse soot. Wolbach et al. (4, 5) and others (7, 8) have assumed that the soot at the K−Pg boundary was caused by global fires ignited by heat from falling spherules. These fires would not have been normal forest fires, but would have acted like large firestorms. Our model is unable to simulate explicitly transport in mesoscale convection, as the horizontal resolution is 2° (∼200 km). Nevertheless, calculations made with mesoscale, nonhydrostatic models indicate rapid lofting of small soot particles to the upper troposphere and even into the stratosphere in intense firestorms (8, 22⇓⇓–25). On the other hand, the coarse particles are more likely to have remained lower in the atmosphere. Thus, we follow the recommendation of Toon et al. (8) and inject the fine soot in a Gaussian near the tropopause and the coarse soot in a half-Gaussian near the surface. Organic carbon would also have been emitted by the fires, but we have no estimates of these emissions. Organics deposited at the K−Pg boundary were likely very different from what is found at present; for example, grasses had barely evolved at the time. Even today, we do not understand which fuels produce brown carbon. Mass fires are much hotter than normal forest fires and may consume the organics in the fire, as suggested by the higher emission rates for soot in intense fires. Organics are also lost by reaction with OH and ozone. The loss rates are poorly known for modern organics, but lifetimes can be on the order of days (55). Thus, while brown carbon would likely have contributed to the short-term aerosol properties, it is not clear that brown carbon would be a significant long-term component of the aerosol lofted into the stratosphere. Because of the uncertainty in the emissions and the added complexity, we do not include organic carbon in our simulations. Pausata et al. (56) did include emissions of 45 Tg of organic carbon and 5 Tg of black carbon in simulations of regional nuclear war, and found that the added organics increases the surface cooling, but that the larger particle size reduces the lifetime of the aerosols, and thus the duration of the climate effects, from 20 y to 10 y. The climate effects in our simulations are of shorter duration and are terminated by an abrupt dehydration event, so they are likely less sensitive to the particle size. The soot in our simulations is subject to both wet and dry deposition. Fig. S2 shows the evolution of the soot burden and cumulative wet and dry deposition for a simulation with a total of 56,000 Tg of soot, of which 41,000 Tg is coarse soot and 15,000 Tg is fine soot. Soot emissions occur from hours 12 to 36, with a sharp increase in both wet and dry deposition occurring near hour 28, 16 h into the fires (Fig. S1A). By the end of the first week, 5.5 d after the end of the fires, the total soot burden is reduced to 50% of the original emission, with 36% being lost to wet deposition and 14% lost to dry deposition. By the end of the first year (Fig. S1B), the soot burden has been reduced to 4% of the total emission, with 57% removed by wet deposition and 38% removed by dry deposition. Assuming the soot that is removed first is the coarse soot near the surface, the lifetime of the coarse soot is 3.4 d, with 75.8% removed by wet deposition, and the lifetime of the fine soot is 5.5 mo, with 60% removed by wet deposition. The lifetime for the coarse soot is similar to an estimate of 3.2 d for the lifetime of black carbon from Jacobson (57). The fine soot has the same lifetime as that for soot injected into the upper troposphere from regional nuclear war calculated by Mills et al. (44) but is shorter than the approximately 1-y lifetime of volcanic aerosols injected into the stratosphere (58). Because the lifetime of the coarse soot is so short, the particle size and composition of these aerosols is unlikely to have a major effect on the overall climate response, so, in a few cases, we omitted the coarse soot entirely.

Aerosol Microphysics Environmental effects from aerosols from a large asteroid impact extend into the mesosphere, and they have important interactions with atmospheric chemistry; therefore, we implemented the CARMA model in WACCM for this experiment. CARMA (12) is a sectional aerosol package, where each aerosol is tracked in a set of size bins. CARMA allows the size distribution of the aerosols to evolve freely, which is necessary when simulating large aerosol injections, as in this study. For the soot aerosol, the fractal treatment of Wolf and Toon (13) is used to better represent the effects of particle aggregation. Optical properties for the soot aggregates are calculated using a mean field approximation (59) assuming a real refractive index of 1.8 and an imaginary refractive index of 0.67. The shortwave optical properties of the fractal soot particles are dependent on monomer size, which is fixed at emission, but are largely independent of particle size, which does increase due to coagulation because of the large soot emissions used in our simulations (60). The Rapid Radiative Transfer Model for GCMs (RRTMG) (61), a radiation package within CESM, is used to include the radiative effects of the impact-generated aerosols. It is likely that, initially, the particles would have a mixed composition of organic and elemental carbon; however, it is hard to know whether one can extrapolate from normal forest fires to what the resulting aerosol would be from a global firestorm. While it would be desirable to treat the aerosol as an internal mixture, knowing the emission of the other aerosols components and being able to treat oxidation of the aerosols is a level of complexity beyond the current capabilities of our model. The optical depth of coated soot particles is about 1.5 times larger than for pure soot particles (62). Thus, our simulations may underestimate the total absorption of the soot particles; however, these particles would also be larger, so their lifetime would be reduced (56).

Model Stability Because the atmospheric response to the impact perturbations is very large, several adjustments to the model tuning and physical parameterizations had to be made to keep the model stable. With these changes, the model remained stable when driven by the exceptionally large chemical and radiative perturbations caused by the impact; we also verified that these modifications had minimal effect on the (unperturbed) control simulation. Changes to Model Dynamics. In WACCM, the dynamical core normally processes a model time step with several smaller dynamics substeps. One phase of the substepping involves a vertical remapping of the model state back to standard hybrid levels. At 1.9° × 2.5° horizontal resolution, WACCM normally uses eight dynamical substeps and performs a vertical remapping every four substeps. For our experiments, we increased the number of dynamical substeps to 128 and carried out the vertical remapping every 8 substeps. In the dynamics, the tendencies on temperature and winds are normally applied fully at the beginning of the dynamical step. Because of the very large temperature and wind tendencies generated by the physical parameterizations in our experiments, we apply a fraction of the tendencies throughout the dynamical step. For temperature, 1/16 of the tendency is applied every vertical remapping step, and, for the winds, 1/128 of the tendency is applied each dynamical substep. Changes to Model Physics. In WACCM, a moist adiabatic adjustment is performed in the troposphere and lower stratosphere, and a dry adiabatic adjustment is performed for the top three layers of the model. Because of the large heating rates, we have extended the dry adjustments to cover all model levels above the level where the moist adjustment stops. The dry adiabatic adjustment parameterization was also modified to mix all tracers in regions of instability rather than just mixing temperature and water vapor. The WACCM chemistry mechanism normally uses explicit calculations for several species that do not change rapidly; however, for these experiments, all of the chemical species are treated implicitly. The RRTMG radiation code was not designed to handle the large temperature and water vapor perturbations generated by our experiments, so some modifications were made to RRTMG to limit extrapolation in lookup tables for optical properties, ensuring nonnegative optical depths and model stability. Because of the large relative humidity produced in the upper atmosphere, the cloud physics parameterization was extended to be active over the entire model domain. The CO 2 injections produced very large mixing ratios of this gas in the upper atmosphere, exceeding the range over which the non-LTE CO 2 cooling parameterization was designed to operate, so we limited the mixing ratio accepted by the parameterization to 700 ppmv. Enhanced cooling in the CO 2 injection case also caused very large temperature gradients in the top layers of the model, so the upper boundary condition for thermospheric temperature was relaxed. Instead of using a temperature from the Mass Spectrometer-Incoherent Scatter (MSIS) empirical model (63) for the upper boundary condition (64), the average of the MSIS and top layer temperatures are used. This choice reduces the temperature gradient sufficiently to prevent failure of the vertical diffusion parameterization.

UVI WACCM does not calculate the UVI, so off-line calculations were performed with the Tropospheric UV and Visible Radiation Model (65), using monthly and zonally averaged output from the CESM simulations and assuming local noon conditions.

Acknowledgments Initial support for C.G.B. and R.R.G. was provided by NASA Grant NNX09AM83G. O.B.T. was supported by the University of Colorado. The authors wish to thank S. Madronich for his comments and for the use of the Tropospheric UV and Visible model. Computational resources supporting this work were provided by both the NASA High-End Computing Program through the NASA Advanced Supercomputing Division at Ames Research Center and by the National Center for Atmospheric Research (NCAR) Wyoming Supercomputing Center, sponsored by the National Science Foundation (NSF) and the State of Wyoming, and supported by NCAR's Computational and Information Systems Laboratory. NCAR is sponsored by NSF.