JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 109, D11309, doi: /2003jd003998, 2004

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 109,, doi: /2003jd003998, 2004 A global climate-chemistry model study of present-day tropospheric chemistry and radiative forcing from changes in tropospheric O 3 since the preindustrial period Sun Wong, 1 Wei-Chyung Wang, 1 Ivar S. A. Isaksen, 2,3 Terje K. Berntsen, 3 and Jostein K. Sundet 2 Received 4 July 2003; revised 15 February 2004; accepted 13 April 2004; published 15 June [1] We present simulations of present-day tropospheric O 3 and of changes in its concentrations and the associated radiative forcing since the industrial revolution. Numerical experiments were conducted using a global tropospheric climate-chemistry model (GCCM) developed by incorporating the University of Oslo (UiO) photochemical module (a reduced tracer scheme that lumps nonmethane volatile organic compounds (NMVOCs) to simulate the CO-NO x -HO x -O 3 system) into State University of New York at Albany Community Climate Model 3 (SUNYA CCM3). The GCCM was run with emissions of pollutants corresponding to two periods, the early 1990s and the preindustrial period, and the simulated tropospheric O 3 is used for off-line calculations of radiative forcing. The study consists of two parts: an evaluation of the GCCM-simulated present-day distributions of tropospheric chemical species with sensitivity experiments to test the effects of NO production by lightning and stratospheric O 3 influx on the simulated results, and the radiative forcing due to O 3 changes since the industrial revolution. The model can reproduce the temporal and spatial variations of O 3 precursors, including CO and NO x (NO+NO 2 ), in the Northern Hemisphere (NH), although there are biases such as the larger CO concentrations in the Southern Hemisphere (SH) and the smaller NO x concentrations in the tropical middle-upper troposphere. The simulated OH concentrations are sensitive to the NO production by lightning. The simulated concentrations of tropospheric OH have a global and annual average of about molecules/cm 3 (corresponding to a tropospheric CH 4 lifetime due to OH of 8.6 years), when the annual production of NO by lightning is about 6 TgN/yr. Compared with O 3 sonde climatology, the model can simulate the seasonal variations of lower to middle tropospheric O 3 within the observed interannual variations when the stratospheric O 3 influx is about 600 Tg/yr, although detailed biases exist in the middle troposphere during fall to winter and in the continental boundary layer during summer. The present-day tropospheric O 3 burden is calculated to be about 376 Tg (or 34 DU) with a net tropospheric chemical production of about 513 Tg/yr. Using realistic stratospheric O 3 concentrations instead of a prescribed O 3 influx as the upper boundary condition enhances the model biases in the extratropical upper-middle tropospheric O 3 concentrations. The model calculates a globally and annually averaged increase in tropospheric O 3 burden of about 13 DU since the preindustrial period, if the production of NO by lightning is kept the same for both the present-day and preindustrial periods. The annually and globally averaged instantaneous radiative forcing due to the O 3 changes is about 0.54 W m 2, with forcing larger than 1Wm 2 occurring over most subtropical to middle-latitude areas in the NH during summer. An assumed 50% reduction in NO production by lightning in the preindustrial period results in an increase of tropospheric O 3 change by 10%, with the associated radiative forcing enhanced by 11%. INDEX TERMS: 0341 Atmospheric Composition and Structure: 1 Atmospheric Sciences Research Center, State University of New York at Albany, Albany, New York, USA. 2 Department of Geosciences, University of Oslo, Oslo, Norway. 3 Center for International Climate and Environmental Research Oslo, Oslo, Norway. Copyright 2004 by the American Geophysical Union /04/2003JD of27

2 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Middle atmosphere constituent transport and chemistry (3334); 0345 Atmospheric Composition and Structure: Pollution urban and regional (0305); 0365 Atmospheric Composition and Structure: Troposphere composition and chemistry; 0368 Atmospheric Composition and Structure: Troposphere constituent transport and chemistry; 1610 Global Change: Atmosphere (0315, 0325); KEYWORDS: tropospheric O 3, radiative forcing, coupled climate-chemistry model Citation: Wong, S., W.-C. Wang, I. S. A. Isaksen, T. K. Berntsen, and J. K. Sundet (2004), A global climate-chemistry model study of present-day tropospheric chemistry and radiative forcing from changes in tropospheric O 3 since the preindustrial period, J. Geophys. Res., 109,, doi: /2003jd Introduction [2] On a worldwide scale, global warming associated with the enhanced greenhouse effect resulting from emissions of anthropogenic CO 2, CH 4, CFCs and N 2 O is considered to lead to significant warming of the earth s surface. The global mean surface air temperature has increased by 0.3 to 0.6 C over the past 100 years, and the observed surface warming is broadly consistent with global climate model (GCM) simulations [Intergovernmental Panel on Climate Change (IPCC), 2001]. Anthropogenic activity can also lead to changes in atmospheric O 3, which is a chemically active radiative species [World Meteorological Organization (WMO), 1999]. Changes in absolute O 3 densities in the lower stratosphere and upper troposphere have been demonstrated to lead to the most pronounced impact on surface temperature in radiative-convective models [Lacis et al., 1990; Wang and Sze, 1980; Wang et al., 1993]. Moreover, two considerations make tropospheric O 3 distinctly different from other greenhouse gases such as H 2 O, CO 2,N 2 O, CH 4, and CFCs. First, O 3 is a secondary constituent whose production involves a large number of chemical processes. A detailed description of the chemical processes is given by Berntsen and Isaksen [1997]. The strong coupling between chemical reactions leads to a pronounced nonlinearity in net O 3 formation with respect to surface pollutant emissions in the troposphere. Crutzen [1979] showed that O 3 formation in the troposphere is highly dependent on NO x levels. As a consequence of the removal of NO 2 through formation of HNO 3 and the subsequent wet removal of HNO 3, O 3 formation in the lower troposphere is estimated to become much less efficient per NO x molecule emitted at high NO x levels than at low NO x levels [Isaksen et al., 1978]. Several studies have demonstrated that O 3 loss takes place in the remote unpolluted troposphere [e.g., Liu et al., 1983], while it is well documented that net O 3 formation occurs over polluted regions. On the basis of observations of key chemical compounds, Ridley et al. [1992] estimated a net O 3 loss of 0.5 ppb/day (1%/day) in the free troposphere near 3.4 km at Mauna Loa. Consequently, the O 3 formation is subjected to large spatial and temporal variations, and the accurate calculation of the formation requires extensive modeling of tropospheric chemistry. [3] The second characteristic that distinguishes tropospheric O 3 from other greenhouse gases is that anthropogenically initiated changes in its distribution show large spatial and temporal variations in the atmosphere because of its short chemical lifetimes in the atmosphere. In the troposphere, O 3 lifetimes range from days to weeks, whereas in the lower stratosphere, they are in the order of months. Therefore radiative forcing from O 3 will be highly nonuniform in space and time, making its impact on climate more difficult to assess than the impact from the well-mixed greenhouse gases. Climate feedback mechanism can further complicate the impact of O 3 increase on the surface temperature. Hansen et al. [1997] demonstrated that O 3 changes in the midtroposphere have the greatest effect on surface temperatures because of cloud feedbacks. Johnson et al. [1999] showed that a warmer climate results in increased concentrations of HO x that mitigate the future enhanced O 3 production caused by the anthropogenic increase of pollutants. Reliable calculations of their climatic effects require the use of models with interactive climate and chemistry. [4] The radiative forcing caused by the future changes in tropospheric O 3 is assessed in many studies (Gauss et al. [2003], Roelofs et al. [1997], Stevenson et al. [1998], etc.). O 3 change since the preindustrial period is also a useful benchmark for the evaluation of model sensitivity, and has been recently estimated by several models to have a substantial impact on the globally and annually averaged radiative forcing [WMO, 1999]. These estimates give values of radiative forcing of 0.35 ± 0.15 W m 2 for tropospheric O 3 increase from the preindustrial period to the present (Berntsen et al. [1997], Mickley et al. [1999], Wang et al. [1998], Shindell et al. [2001a], etc.), with stratospheric adjustment considered in most of the estimates. These studies further indicate that tropospheric O 3 increases since the preindustrial period could augment the radiative forcing from all other greenhouse gases by about 10% to 20% over the same period. More importantly, there are strong regional variations in the radiative forcing from the O 3 changes. Regional O 3 changes in areas where there have been strong increases in the emissions of pollutants over the last couple of decades (e.g., Southeast Asia) are calculated to have caused radiative forcings in some regions of up to 0.5 W m 2 [Berntsen et al., 1996; Fuglestvedt et al., 1999]. [5] In addition to the spatial variation in the radiative forcing, a study by Mickley et al. [2001] showed that the instantaneous radiative forcing of O 3 changes since the preindustrial period is subjected to uncertainties in the input parameters for the calculation of the chemistry in the preindustrial times. After adjusting the emissions of NO x from lightning, soils, and the emissions of biogenic hydrocarbons, the instantaneous radiative forcing can be enhanced to values of W m 2. It should be noted that these instantaneous forcings should not be directly compared with the range of 0.35 ± 0.15 W m 2 quoted above, which mainly describes the uncertainty range of the forcings with stratospheric adjustment. Nevertheless, the increase of instantaneous forcing reported by Mickley et al. [2001] is significant compared with the instantaneous forcing of 0.44 W m 2 [Mickley et al., 1999], which is the control forcing computed by the same model without any 2of27

3 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 adjustments of the input parameters. The reduction in the preindustrial lightning source accounts for two thirds of the increase in instantaneous forcing [Mickley et al., 2001]. Estimates of NO production from lightning cover a wide range of uncertainty and are highly related to the convective activity in the troposphere. Model simulations of the doubled CO 2 climate show intensified penetrative convection in a future warmer climate [Rind et al., 2002]. Price and Rind [1994] also implied an increase (decrease) in global lightning frequencies in a warmer (cooler) climate. This result may imply a smaller NO production rate from lightning in the cooler preindustrial period. [6] A coupled climate-chemistry model is developed and evaluated in this study with available observations. The model contains a chemical module that includes the oxidation processes of nonmethane hydrocarbons (NMHCs) [Berntsen and Isaksen, 1997; Sundet, 1997] that are necessary for the calculation of the nonlinear tropospheric O 3 formation. The model also contains a simple stratosphere (with coarse vertical resolution) that can generate reasonable climate variability in the upper troposphere and lower stratosphere [Wong and Wang, 2000, 2003]. Shindell et al. [2001b] demonstrated that proper treatments of dynamical stratosphere-troposphere interaction are necessary for simulations of future climate. The model is then applied to conduct sensitivity experiments to test the effects of the uncertainty in the NO production of lightning on the present-day tropospheric chemistry. The uncertainty in the estimated radiative forcing is also estimated by conducting a sensitivity run with a smaller NO production rate from lightning in the cooler preindustrial period. Section 2 will describe the model structure in more detail, and section 3 will evaluate the model behavior with available observations. In section 4, simulations of the seasonal variation of tropospheric O 3 as well as the O 3 budget will be illustrated. The model is applied in section 5 to estimate the radiative forcing of O 3 changes since the preindustrial period. Conclusions and discussion are included in section Model Structure 2.1. General Circulation Model [7] The global tropospheric climate-chemistry model (SUNYA/UiO GCCM) is developed by incorporating a simplified version of the UiO tropospheric photochemical scheme into SUNYA CCM3. The SUNYA CCM3 is based on National Center of Atmospheric Research Community Climate Model version 3 (NCAR CCM3) [Kiehl et al., 1998] with SUNYA four-dimensional (4-D) O 3 [Wang et al., 1995; Wong and Wang, 2000, 2003] and the sea surface temperatures (SSTs) supplied for Atmospheric Model Intercomparison Project 2 (AMIP 2) [Gates et al., 1999] as inputs. The GCCM uses T42 horizontal resolution (equivalent to ) with 18 layers in the vertical extending from the surface to about 2.5 hpa (2 km vertical resolution near the tropopause). The dynamical time step is 20 min, while shortwave and longwave fluxes are calculated every hour. The model can provide reasonable climatology of present-day climate parameters for use in simulations of tropospheric chemistry [Hack et al., 1998; Hurrell et al., 1998; Kiehl et al., 1998], although there are cold biases of about 4 12 K in polar regions and small dry biases in summertime NH middle latitudes. The model can also reproduce reasonable interannual variability of the tropospheric system (e.g., ENSO and tropopause temperature [Hack et al., 1998; Wong and Wang, 2003]). [8] For tracer transport, we replace the original built-in semi-lagrangian (SLT) scheme in CCM3 with a flux-form semi-lagrangian scheme, the Split Implementation of Transport Using Flux Integral Representation (SPITFIRE) [Rasch and Lawrence, 1998]. Although the original SLT scheme is efficient and economic for tracer transport in global climate models, it suffers from two major problems: (1) its inherently nonconservative nature; and (2) poor performance in a coarse vertical resolution model in the vicinity of the tropopause, where the atmosphere makes a transition from a relatively well-mixed turbulent fluid to a stably stratified one. The SPITFIRE scheme incorporates a flux form scheme into the SLT scheme to ensure the conservative property. Tests of the scheme for an O 3 -like tracer [Rasch and Lawrence, 1998] indicate that at 100 hpa, SPITFIRE converges faster than SLT for an 18-layer model. Moreover, the result of SPITFIRE at 100 hpa from the 18-layer model resembles that of a 46-layer model with higher resolution around the tropopause, although the results from the two models differ at 65 hpa [Rasch and Lawrence, 1998] Chemistry [9] A state-of-the-art photochemical scheme used in global chemical tracer models (CTMs) typically includes chemical species, of which at least 50% have lifetimes long enough for transport to significantly influence their distributions. Integrating a photochemical scheme of this nature into a global climate model increases the CPU time needed for the calculations and the requirements for computer memory. For this reason, the total number of species (both transported and in local steady states) should be kept at a minimum. Two approaches were taken to simplify the comprehensive chemical scheme [Isaksen and Hov, 1987; Berntsen and Isaksen, 1997, 1999], which involves approximately 50 components with about 120 thermal reactions and 18 photolysis reactions, and make it more efficient for use in a global climate model. First, we consider the number of NMHCs needed in the model. Second, we consider the possibility of treating the semishort-lived species as being in local photochemical steady states, thereby reducing the number of species to be transported in the model. [10] The first approach is based on the idea that the number of NMHCs can be reduced without a significant loss of accuracy when the chemical scheme is designed for studies of perturbations of the free troposphere. Some of the NMHCs are oxidized rapidly by OH in the planetary boundary layer (PBL), so that they mainly have an indirect effect on the chemistry of the free troposphere. In the simplified scheme, we have chosen to represent this class of NMHCs (C 2 H 4,C 3 H 6 and m-xylene from the original scheme) by adding their emissions to the emissions of isoprene (for more detailed explanation see Table 1). Similarly, emissions of the higher-order alkanes (butane and hexane) are lumped together with and treated as propane. This approach allows us to reduce the number of species in the scheme by 17, including the oxygenated and radical 3of27

4 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Table 1. Chemical Species Used in the Original and Simplified Schemes Components treated equally in both schemes Components not transported in the simplified scheme Components lumped with other components in the simplified scheme Components omitted in the simplified scheme Components O 3, HNO 3, PANX, a CO, CH 4, b C 2 H 6,C 3 H 8, HCHO, CH 3 CHO, H 2 O 2,HO 2 NO 2, isoprene, ISOK, c acetone, NO, NO 2,NO 3,N 2 O 5, O3NO, d,e NOX, d,f NOZ, d,g O( 3 P), d O( 1 D), d OH d HO 2,CH 3 O 2,C 2 H 5 O 2,C 3 H 7 O 2, ISOR1, h ISOR2, i CH 3 COO 2 C 2 H 4, j C 3 H 6, j C 4 H 10, k C 6 H 14, k m-xylene j C 4 H 9 O 2,C 6 H 13 O 2,CH 3 COCH(O 2 )CH 3, CH 3 CH(O 2 )CH 2 OH, AR1, l AR2, m AR3, n HCOHCO, RCOHCO, CH 3 COCOCH 3, CH 3 COC 2 H 5,CH 3 COCH 2 (O 2 ) a PANX = PAN + CH 3 COO 2. b CH 4, fixed in this study. c ISOK, sum of methylvinylketone and methacrolein. d NT, species not transported in both schemes. e O3NO = O 3 -NO, surrogate component used to avoid the effect of the rapid cycle O 3 +NO! O 2 +NO 2 ;NO 2 +hn! NO+O;O+O 2 +M! O 3 +M [cf. Berntsen and Isaksen, 1997] for details. f NOX = NO + NO 2 +PAN+NO 3 +HO 2 NO 2 +2N 2 O 5. g NOZ = NO 3 +N 2 O 5. h ISOR1, RO 2 radicals produced through the reaction isoprene + OH + O 2. i ISOR2, RO 2 radicals produced through the reaction ISOK + OH; see Berntsen and Isaksen [1997] for details on the isoprene chemistry. j Lumped with the emission of isoprene. k Lumped with the emission of C 3 H 8. l AR1, RO 2 radical produced through the reaction m-xylene + OH + O 2. m AR2, A C-5 carbonyl compound formed from AR1. n AR3, A C-5 RO 2 radical formed from AR2. products in the oxidation chains. Generally, the full oxidation chain, with reaction products as in the comprehensive scheme [Berntsen and Isaksen, 1997], has been kept unchanged for those NMHCs that are represented in the simplified scheme. The only exceptions are in the reactions of ISOR2 with NO or CH 3 O 2 (see Table 1), where the production of RCOHCO is replaced by an equal production of CH 3 CHO, and in the reaction of OH + acetone, where the production of CH 3 COCH 2 (O 2 ) is replaced by an equal production of CH 3 CO. [11] The second approach has been implemented by generally treating all peroxy radicals (HO 2 and RO 2 )as being in local photochemical steady states, thus omitting the transport of these species. This reduces the number of transported species by 7 (many of the original peroxy radicals in the scheme were omitted through the reduction of the number of NMHCs). Altogether, the simplified scheme consists of 51 species, of which 18 are transported. In this paper, the CH 4 concentration is fixed at 1.76 ppmv for the present-day climate and 0.7 ppmv in the preindustrial period, and hence the number of tracers is reduced to 17. [12] A 3-D global chemistry transport model, Oslo-CTM2 [Sundet, 1997], was used to test the performance of the simplified scheme relative to the comprehensive scheme. After some initial tests and revisions in a simple box model, the simplified scheme was introduced into the global CTM. Two control simulations of 7 months, using the same initial conditions, emissions and meteorology, were performed with both the comprehensive scheme and the simplified scheme. The simulation with the simplified scheme includes the reduction in the number of transported species as given in Table 1. Since we are primarily interested in radiative effects of perturbations in the chemical composition from human-made emissions, we have performed perturbation studies with the full chemical scheme and the simplified scheme and compared these studies with the control runs. In the perturbation studies we increase the yearly NO x emissions from anthropogenic sources over Southeast Asia by 1 Tg (N). Figure 1 shows the difference between the control and the perturbed cases of the calculated zonally averaged July O 3 concentrations for the simulations with the extensive chemical scheme and those with the simplified scheme. We have chosen to perform the perturbation comparisons at these relative low latitudes (compared to perturbations over Europe or the United States) since the photochemistry is more active at lower latitudes. The NO x emissions at low latitudes should lead to more significant chemical perturbations compared to those at higher latitudes. We see that the patterns in the O 3 perturbations are rather similar with regard to both vertical and latitudinal distributions, with slightly more increases in O 3 formation when we use the simplified scheme than when we use the extensive chemical scheme. [13] There are still significant uncertainties in our understanding of the chemistry of the troposphere and in the quantification of major sources of O 3 precursors (e.g., NO x from lightning, isoprene from deciduous forests). Given these uncertainties and the fact that O 3 changes in the free troposphere have a larger impact on radiative forcing than changes in the boundary layer, we believe that the accuracy of the simplified chemical scheme is acceptable for simulations of radiative effects of tropospheric O 3 changes in a GCM. [14] Photodissociation coefficients (J values) are calculated by a fast J value scheme [Wild et al., 2000], which has an online treatment of absorption and scattering by both molecules and aerosols. The J values are updated every 4of27

5 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 are used. Over continents, the ice/snow values are used if the surface temperature is below 3 C; the land values are used if the surface temperature is over 3 C; and the linear interpolated values between the ice/snow values and the land values are used if the temperature is between 3 C and 3 C. If the surface temperature is over 35 C, we assume that the case is over a desert and the ice/snow values are used. [16] In the model, soluble species (O 3, HNO 3, HCHO, H 2 O 2,CH 3 OOH, and ISOK (sum of methylvinylketone and methacrolein; see Table 1)) will undergo wet deposition. The scavenging scheme represents a first-order process similar to that used by Brasseur et al. [1998]. dc i dt ¼ bc Figure 1. Calculated increase in zonally averaged ozone (ppbv) in July due to 1 Tg(N)/yr of extra NO x from fossil fuel sources in SE Asia. The upper panel is the result from the original chemical scheme, and the lower panel is from the simplified chemical scheme. hour according to the instantaneously calculated surface albedo and cloud optical thickness. The O 3 and temperature used in the computation are prescribed by off-line climatologies, in order to avoid errors caused by the model biases in the simulated O 3 and temperature. [15] Heterogeneous removal of N 2 O 5 on aerosols, which is likely to be important for tropospheric O 3 at the NH high latitudes during the winter months, is incorporated by applying the zonally and annually averaged k coefficients given by Dentener and Crutzen [1993]. Loss of species from dry deposition is proportional to the deposition velocities at 1 m above the surface, which are listed in Table 2, and the local concentrations of the corresponding gases. Because of the coarse vertical resolution in the boundary layer, the concentrations of the gases have to be extrapolated to the surface using the online calculated surface heat diffusivity [Berntsen and Isaksen, 1997; Isaksen and Rodhe, 1978]: dc L dt C L ¼ V 0 z L ð1 þ V 0 z L =KÞ where C L is the gas concentration at the model s lowest layer; K is the model s surface heat diffusivity; z L is the height of the midpoint of the model s lowest layer; and V 0 is the deposition velocity at 1 m above the surface. Over the ocean or sea ice, the corresponding sea or ice/snow values where C i is the species concentration at the ith level, C is the species concentration at the ith level if rain formation occurs, and at the (i 1)-th level (one level above) if reevaporation occurs; and b is the loss coefficient (s 1 ). The efficiency of the uptake in the droplets is a function of the solubility of the gases in the aqueous phase (i.e., the temperature-dependent Henry s constant). The loss coefficient is proportional to the rain formation rate and cloud fraction as follows: b ¼ dq=dt F q av 1 1 þ 1= ðhrtq cld Þ where dq/dt is the model rainwater tendency (g cm 3 s 1, which is positive if formation of rain occurs and negative if re-evaporation occurs); q av is the available precipitation defined as the ratio of cloud water density to the cloud fraction (q av = q cld /F); q cld is the cloud water density (g cm 3, including both liquid and ice cloud); F is the cloud fraction; H is the temperature-dependent Henry s constant (M atm 1 ); R is the ideal gas constant ( atm cm 3 K 1 M 1 g 1 ); and T is the temperature (K). We distinguish whether a species is removed because of rain formation, or released back to its gas phase because of re-evaporation. If the local dq/dt is positive, removal of species occurs, and every variable on the right-hand side is taken as having the local value. If the local dq/dt is negative, re-evaporation occurs, and every variable (except dq/dt) on the right-hand side is taken as having the value of one layer above (i.e., the (i 1)-th level). For the highly soluble HNO 3, we assume its Henry coefficient is very large and the value in the bracket then becomes equal to Emissions [17] The surface emissions of biogenic and anthropogenic species are based on the data suggested by IPCC [2001] Table 2. One-Meter Dry Deposition Velocities Used in GCCM a Species Land Sea Ice/Snow O NO HNO PAN H 2 O CO a Velocities are given in cm/s. 5of27

6 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Table 3. Global Emissions of Pollutants in GCCM a Species Industrial Biomass Burning Biogenic and Oceans Aviation Lightning Total CO b C 2 H C 2 H C 3 H C 4 H C 6 H m-xylene C 3 H CH 3 COCH Isoprene c NO c c NO a Emissions are given in Tg/yr. b Ocean emission of CO is reduced to 25 Tg/yr from 50 Tg/yr. c Unit for isoprene is TgC/yr, and units for NO and NO 2 are TgN/yr. (OxComp Y2000). The annual mean values are rescaled so that the emission rates are more suitable for model evaluations using observations in the early 90s and comparable to those used in other CTMs or photochemical GCMs. Table 3 lists the globally and annually averaged surface emission rates for the species used in the model. [18] For CO, biogenic emission includes emissions from vegetation and oceans, and is about 175 TgCO/yr, with the ocean emission reduced from the value of 50 TgCO/yr to a value of 25 TgCO/yr. Biomass burning emission includes emissions from fires in forest, savanna, deforestation, and waste burning, and is about 515 TgCO/yr. Industrial emission includes emissions from combustion of fossil and domestic fuels, and is about 477 TgCO/yr. Both biogenic and industrial emissions are assumed to be uniformly distributed in a year. The seasonal variation of the emission from biomass burning at each grid is scaled according to that of the CO 2 emission based on Müller [1992]. [19] The surface emission of NO x includes emissions from natural soil, biomass burning, and fossil fuel combustion. Global Emissions Inventory Activity (GEIA) data are used for the distribution of soil emission, with the total emission scaled to the value of 5.5 TgN/yr. The emission from biomass burning has a seasonal variation as estimated by Müller [1992], with the total emission scaled to the value of 6.4 TgN/yr. The emission from fossil fuel combustion is scaled to the value of 27.1 TgN/yr. All NO x surface emissions are assumed to consist of 97% NO and 3% NO 2. The injection of NO from aviation is distributed three dimensionally according to a NASA report [Friedl et al., 1997] and assumed to have no seasonal variation, corresponding to a total rate of about 0.7 TgN/yr. Production of NO by lightning is parameterized according to Price et al. [1997]. Monthly climatology of NO production is specified for each latitude band and scaled to an annual emission rate of 3 TgN/yr. The production of NO is distributed in the longitudinal direction by a factor proportional to the local convective mass flux in the updraft and a power of the height of the top of convection activity (H 4.92 over continent and H 1.73 over ocean). Then, the produced NO molecules are vertically distributed by the prescribed altitude-dependent fractions given by Pickering et al. [1998]. [20] The IPCC [2001, their Table 4.7b] partitions as well as the total emission rates for the industrial (210 Tg/yr) and biomass burning (42 Tg/yr) sectors were used to obtain emissions of the NMVOCs used in the model. The seasonal variation of the biomass burning is again scaled according to the CO 2 emission based on Müller [1992]. The total emission rates for the NMVOCs used in the model as well as the partitions for industrial and biomass burning sectors are listed in Table 3. Isoprene emission from tree foliage is based on the GEIA project with seasonal variations. The annual emission is rescaled to the value of 220 TgC/yr as suggested by the OxComp Y2000 scenario [IPCC, 2001] Stratosphere-Troposphere Exchange [21] In the GCCM, transport calculations are carried out in the whole domain of the model, while the chemistry computations are carried out only in the troposphere, starting at the fifth layer from the top of the model (about 99 hpa). Previous model studies demonstrated that strato- Table 4. Regions and Periods of the Aircraft Data Sets Used in This Study Region Name Latitude Range Longitude Range Period and Campaign E. Japan N Feb. to 14 March 1994 (PEM-West-B) E. China N Feb. to 14 March 1994 (PEM-West-B) 16 Sept. to 21 Oct (PEM-West-A) Philippine 5 20 N Feb. to 14 March 1994 (PEM-West-B) Tahiti 20 S Aug. to 5 Oct (PEM-Tropics) E. Brazil 15 5 S Sept. to 26 Oct (TRACE-A) S. Africa 25 5 S Sept. to 26 Oct (TRACE-A) W. Africa 25 5 S Sept. to 26 Oct (TRACE-A) 6of27

7 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 sphere-troposphere exchange of O 3 plays an important role in the O 3 budget in the upper troposphere [Lelieveld and Dentener, 2000], and the net O 3 production rate in this region is sensitive to the local NO x concentration [Friedl et al., 1997]. A proper treatment for the computation of stratospheric O 3,NO x, and HNO 3 is necessary to account for the stratosphere-troposphere exchange of these species. In GCCM, synthetic O 3 (Synoz) [McLinden et al., 2000] is used as the stratospheric O 3 above 80 hpa to account for the upper O 3 boundary condition. A fictitious stratospheric O 3 production, which matches the annual mean stratospheric O 3 influx rate of about 600 Tg/yr, is uniformly distributed in the tropics between 7 and 81 hpa (the second to fourth model layers). The model dynamics will then determine the spatial and temporal variations of the O 3 flux across the tropopause. As a sensitivity study, we have performed an experiment using another treatment in which the stratospheric O 3 concentrations (e.g., above 99 hpa) are prescribed according to values derived from observations (SUNYA O 3 climatology in our case). The O 3 flux across the 99-hPa level is then determined by the model air mass flux and the O 3 concentrations at that level. The sensitivity of tropospheric chemistry to the treatments of stratospheretroposphere O 3 exchange can then be tested. [22] For NO x and HNO 3, a simple correlation of NO y to O 3 mixing ratios is specified in the lower stratosphere, and the partition between NO x and HNO 3 is assumed to be 1:3 in the lower stratosphere. The correlation of NO y to O 3 mixing ratios is chosen to be , falling within the range ( to ) given by observations [Fahey et al., 1996; Murphy et al., 1993]. 3. Model Evaluation [23] The control simulation was forced by the climatological SSTs for 3.5 years using the synthetic O 3 for the upper boundary condition. The initialization used the species fields obtained from Oslo-CTM2 and started in July. The model climatologies of each species for each month are computed from the last 3 years of the control simulation. The model interannual variation of a species concentration is computed as the standard deviation of the concentration from the 3-year results. On the basis of the control experiment, two sensitivity experiments were also performed for 18 months, and the results of the last 12 months were used for analyses. In the first experiment (Exp-flux), the stratosphere-troposphere O 3 exchange is computed by the prescribed stratospheric O 3 concentrations above 99 hpa according to SUNYA O 3 climatology. The stratospheric O 3 influx in Exp-flux is calculated to be about 912 Tg/yr, about 52% larger than the flux in the control experiment. Accordingly, the NO x and HNO 3 fluxes from the stratosphere were increased from the control experiment values. In the second experiment (Exp-ligtn), the total NO production by lightning was raised to the value of 6 Tg/yr from the value of 3 Tg/yr. In this section, the control simulation of O 3 precursors is evaluated with available observation data sets, while the results of the sensitivity simulations will be mentioned to diagnose the origins of the model biases. [24] The observation data sets include monthly mean surface CO mixing ratios provided by National Oceanic and Atmospheric Administration, Climate Monitoring and Diagnostics Laboratory (NOAA/CMDL) [Novelli et al., 1998], the composites of regional aircraft measured profiles for O 3 precursors provided by Emmons et al. [2000], and the ozonesonde climatology compiled by Logan [1999]. For comparisons with observations over a site (e.g., the NOAA/CMDL surface CO or ozonesonde climatology), the monthly mean model species concentrations of a grid closest to the site were chosen. For comparisons with the regional aircraft profile composites, the monthly mean model data of the grids within the region and of the months when the campaigns took place were averaged. Table 4 shows the latitudinal and longitudinal ranges of the regions in the NASA/GTE aircraft campaigns that were chosen for our model evaluation. Detailed geographical locations of the regions are given by Emmons et al. [2000]. For most of the species discussed in this section, we chose the profiles in which at least 5 altitudes have averages that use more than 15 measurements to ensure better statistics. For HNO 3, the measurement number criterion is reduced to 10 from 15 for the sparsely available HNO 3 data Carbon Monoxide [25] CO is mainly emitted by fossil fuel combustion and biomass burning. It can also be produced by oxidation of other VOCs (e.g., CH 4 and isoprene). Most of the removal of CO is through its oxidation by OH. The seasonal variations of surface CO concentrations from the control and Exp-ligtn simulations are plotted against the NOAA/CMDL climatologies over a few sites in Figure 2. In general, the model can satisfactorily reproduce the seasonal cycles and the magnitudes of the concentrations over the sites in the NH. However, in the SH, the model overestimates CO concentrations in both control and Exp-ligtn simulations for all seasons. Another sensitivity test (not shown) was carried out in which the ocean emission was shut off. The CO concentrations in the SH were still larger than the observed values. The overestimation is, therefore, caused by some other factors. The following section shows that the annually and globally averaged OH concentrations in the control and Exp-ligtn simulations span the range of molecules/cm 3 (corresponding to tropospheric CH 4 lifetimes due to OH of about years), while the corresponding changes in CO concentrations at the SH surface between the two simulations are insignificant compared to the model interannual variability. Thus the overestimated SH CO concentrations are partly related to our smaller OH concentrations (see section 3.3). The CO biases in the SH over regions affected by biomass burning (Ascension Island) are reduced during the biomass burning season. Over Cape Grim, which is a remote region, the bias is less dependent on the season. Given the above facts, we speculate that the SH CO bias is also partly related to some dynamical reasons instead of biases in the surface emissions. Hauglustaine et al. [1998] presented the results of a CTM (model for ozone and related chemical tracers, MOZART), which used the CCM2 W0.5 winds to transport tracers (an intermediate version between CCM2 and CCM3), and showed overestimation of surface CO concentrations over Samoa and Cape Grim during January to March. Another CTM 7of27

8 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 2. Seasonal variations of surface CO concentrations (in ppbv) for the 3 control model years (thin solid and dashed lines) and for Exp-ligtn (thick dashed lines, for definition see the footnote of Table 5) compared with the CMDL climatology (thick solid lines). The vertical bars show the interannual variations (standard deviation of different years results) of the CMDL climatology. (MATCH-MPIC 2.0 [Lawrence et al., 1999]) used the same transport scheme (SPITFIRE) but reanalysis winds (NCEP) to transport tracers, and it also overestimates CO in the SH high latitudes (Syowa), which is relatively remote from the anthropogenic influence. Over Ascension Island, the model simulates a peak in August, about one month earlier than the month of the observed peak, which is caused by biomass burning. The interannual variability of the model CO is generally smaller than the observed variability, probably owing to the lack of interannual variations in the emissions of biomass burning and in the SSTs that can force the interannual variation in the surface circulation. [26] Figure 3 compares the vertical profiles of CO from both control and Exp-ligtn simulations with the regional aircraft profiles over some regions. The model satisfactorily reproduces the vertical profiles over the polluted area (coast of China) and its corresponding outflow area (east coast of Japan in February and March). Although the vertical variation is well simulated, the model slightly overestimates the CO concentrations over the Philippine Sea in February March, and the overestimation is more evident over Tahiti, which is in the SH. This is consistent with the results shown in Figure 2 that the model tends to overestimate CO concentrations in the SH. The simulations of biomass burning over South Africa and of the corresponding outflow area over the west coast of Africa are reasonable, although the altitude and magnitude of the peak over the west coast of Africa are both underestimated. Over east Brazil, the model seems to have a smaller source of biomass burning as well as the associated CO concentrations in the troposphere. Increasing NO production by lightning results in more conversion of HO 2 to OH in the tropics and, consequently, slightly reduces the CO concentrations in the SH troposphere Nitrogen Compounds [27] In the tropics, major sources of NO x are biomass burning during dry seasons as well as lightning in the upper troposphere. In the extratropics, NO x is directly emitted through fossil fuel burning over industrialized areas. Loss of NO x occurs mainly through NO 2 and OH reacting to form 8of27

9 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 3. Model vertical profiles of CO concentrations (thick solid lines for the control and thick dashed lines for Exp-ligtn, see the footnote of Table 5 for definition) compared with the composites of aircraft measurements [Emmons et al., 2000] (thin solid lines as the median and dashed lines as 25th and 75th percentiles). The period and region of the data are shown in the title of each figure. Detailed locations of the regions are included in Table 4. HNO 3. During winter nights when temperatures are low, NO x can also form NO 3 and N 2 O 5, which are removed by heterogeneous reactions. Figure 4 shows the vertical profiles of the simulated NO x concentrations along with the concentrations derived from the aircraft NO data sets. Over polluted areas the model can simulate the C-shape profiles with higher concentrations in the lower and upper troposphere. In the extratropics over the east coast of China, where industrial surface emissions are high, the model reproduces reasonable NO x profiles. Over biomass burning areas in the tropics, the model underestimates NO x concentrations in the middle-upper troposphere. The reason for this is unclear since the local surface levels of NO x are not low, and convective transport in the model is efficient. Increasing the production of NO by lightning results in increased upper tropospheric NO x concentrations; however, this increase is seasonally dependent and does not account for the discrepancy between the model and the observed results. Since the underestimation of NO x levels is more evident over the convectively active areas in the tropics (e.g., the Philippine Sea and South Africa), it is possible that the detailed parameterization of NO production by lightning should be improved in this model. For example, most of the increased NO x concentrations are located in altitudes above 11 km in 9of27

10 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 4. Similar to Figure 4 but for NO x vertical profiles (in pptv). the model, whereas the peaks of the observed NO x are located in lower altitudes, suggesting that improvement in the vertical distributions of the produced NO x is necessary. [28] HNO 3 concentrations over South Africa and the Philippine Sea (Figure 5) are somewhat underestimated, consistent with the local underestimation of NO x concentrations. Over the west coast of Africa the HNO 3 profile is well reproduced. In this area there is a relatively small surface emission of NO x and the local vertical downward motion can bring in stratospheric NO y to the upper troposphere. The lower tropospheric HNO 3 over this area (1 3 km) is transported from South Africa by the easterlies. However, since the NO x and HNO 3 concentrations over South Africa are underestimated, it is likely that the good match with observations over this site is just a coincidence. The underestimated upper tropospheric NO x in this area probably implies that a better treatment for the tropical stratospheric-tropospheric exchange of NO y is necessary (in this model the O 3 /NO y ratio in the lower stratosphere is assumed to be spatially uniform) Hydrogen Compounds [29] In the troposphere, the production of OH is mainly governed by the following reactions: H 2 O þ O 1 D! 2OH NO þ HO 2! OH þ NO 2 The first reaction is important for the OH production in the moist lower troposphere, while the second is important in the upper troposphere, especially over areas where the reaction is not saturated. OH is converted back to HO 2 10 of 27

11 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 computed the lifetime of methylchloroform as the inverse of the air-mass-weighted average of the product k(t) [OH] in the troposphere, where k(t) = e 1550/T [DeMore et al., 1994]. The results are 5.5 and 5.1 years for the control and Exp-ligtn experiments, respectively, both longer than the derived value of 4.9 ± 0.3 years for the lower atmospheric lifetime due to OH by Prinn et al. [1995]. We also compute the tropospheric lifetime of CH 4 using the same method with k(t) = e 1800/T. The results are 9.1 years and 8.6 years for the control and Exp-ligtn experiments, respectively. While Prinn et al. [1995] gave a value of 8.0 ± 0.5 years for the CH 4 lower atmospheric lifetime due to OH, Krol et al. [1998] computed a value of 8.6 years for 1993 by claiming an upward trend in OH of 0.4% per year and IPCC [2001] suggested a value of 9.6 years. [30] H 2 O 2 is formed by the reaction of two HO 2 radicals and removed by wet deposition and photodissociation. Investigation of the profiles of H 2 O 2 is an indirect way to evaluate the model simulated HO 2, which is essential for tropospheric O 3 production in polluted areas and loss in remote areas. Figure 7 shows the model vertical profiles of H 2 O 2 along with the aircraft measured profiles. In general, the model can satisfactorily simulate profiles of H 2 O 2 in the chosen regions. Figure 5. Similar to Figure 4 but for HNO 3 vertical profiles (in pptv). through reactions with CH 4, CO, and NMVOCs. Figure 6 shows the monthly and zonally averaged OH concentrations for January and July from both control and Exp-ligtn runs. The distributions are similar to those shown by Spivakovsky et al. [2000] with the maxima of OH concentrations located around hpa. Increasing NO production by lightning from 3 Tg/yr to 6 Tg/yr can increase the maxima of the OH concentrations. The annual and global average of the OH concentrations is molecules/cm 3 for the control experiment and molecules/cm 3 for Expligtn experiment, lower than the derived value of molecules/cm 3 by Spivakovsky et al. [2000]. The smaller OH concentrations in the model are related to the overestimation of CO in the SH and the underestimation of NO x in the tropical middle-upper troposphere. We 4. Seasonal Variations of Tropospheric O 3 and Its Budget [31] Both dynamical and chemical processes control the budget of tropospheric O 3. As different processes undergo different seasonality and dominate in different spatial regions, tropospheric O 3 has large seasonal variations that also largely depend on the regions of interests. The difficulty and uncertainty in simulations of different processes also result in a wide range of model O 3 budgets in the literature [Shindell et al., 2001a]. In this section, we evaluate our model simulations of tropospheric O 3 by comparing the model results with the O 3 sonde climatology [Logan, 1999]. Since changes in NO production by lightning have relatively small effects on temporal and spatial O 3 variations compared to changes in treatment of stratospheric O 3 influx over the sonde sites being shown, the evaluation uses only the results of the 3-year control and the 1-year Exp-flux experiments Evaluations With O 3 Sonde [32] Figure 8 shows the seasonal variations of O 3 over several sonde sites from the sonde climatology and the model results for 300, 500, and 800 hpa. In the lower troposphere (800 hpa), the model can reproduce reasonable seasonal variations at different sites in both hemispheres. The model seasonal variation in the NH middle latitudes are reasonable compared with the observations over western and eastern United States (Edmonton, Boulder, and Wallops Island), central Europe (Hohenpeissenberg), and eastern Asia (Sapporo), with winter minima and summer maxima of O 3 concentrations. Detailed model departures from the observations in these locations include the slightly larger O 3 concentrations during fall to spring and the overestimation during summer. During fall to spring, the overestimation is enhanced for the results 11 of 27

12 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 6. Model zonal mean OH (in 10 5 molecules/cm 3 ) distributions for (left) January and (right) July. The upper panel is for the control climatologies, and the lower panel is for Exp-ligtn (for definition, see the footnote of Table 5). of Exp-flux, indicating that downward transport of stratospheric O 3 plays a role in causing model biases in midlatitude regions. In summer, dry deposition is important because of the elevated boundary layer height. The overestimated summertime O 3 concentrations over these sites imply that the modeled dry deposition rates for O 3 are too weak in these regions. Over Kagoshima, which is affected by the Asian outflow of pollutants during late winter to spring and the ventilation during summer monsoon season, the model can capture the O 3 seasonal variation with the spring peak and the summer dip. However, the overestimations of O 3 concentrations over Sapporo and Kagoshima exceed the observed interannual variability in winter. Over Hilo, a relatively remote area, the model can simulate the spring to early summer O 3 but calculates larger O 3 concentrations in late summer to winter. The effect of the larger stratospheric O 3 influx is not evident, consistent with the fact that Hilo is located in the tropics where upward motion is more dominant. In the SH, the model reproduces seasonal variations of O 3 at 800 hpa reasonably well. Detailed departures include the smaller O 3 concentrations over Natal during September November and the overestimation over Lauder during April June. The smaller O 3 concentrations over Natal are consistent with the underestimation of CO and NO x concentrations over eastern Brazil at the same altitude (sections 3.2 and 3.3), implying too weak emissions from biomass burning in the model in this area. The enhancement of the overestimation over Lauder in Exp-flux suggests that the too efficient downward transport may have some contributions to the local model biases. [33] In the middle troposphere (500 hpa), the seasonal variations of O 3 are very similar to those at 800 hpa with more fall-to-spring biases that result in the smaller model seasonal variations in the NH extratropics. Over Boulder and Wallops Island, the model O 3 concentrations have peaks later than the observed peaks. Larger stratospheric O 3 influx has evident effects in the NH middle latitudes from September to March, suggesting that the biases are mainly due to the efficient downward transport of the stratospheric O 3 caused by the SPITFIRE scheme (P. Rasch, personal communication, 2003). MATCH-MPIC 2.0 [Lawrence et al., 1999] used the SPITFIRE transport scheme and also simulated larger tropospheric O 3 concentrations in winter to spring over the NH high latitudes. In our model, the biases become more evident in fall to early spring over most of the sites because of the relatively longer O 3 lifetime in this season. Over Hilo, the model O 3 seasonality is also smaller than the observed seasonality, and the concentrations are overestimated from July to November. In the model the 12 of 27

13 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 7. (in pptv). Similar to Figure 4 but for H 2 O 2 vertical profiles smaller biomass burning activity over Natal and the larger O 3 concentrations in April May over Lauder become more evident at 500 hpa than at 800 hpa. [34] In the NH upper troposphere (300 hpa), the phases of the observed seasonal O 3 variations in middle latitudes begin to differ from those in the stratosphere. The control experiment can reasonably reproduce the O 3 seasonal variations over middle latitudes (Edmonton, Boulder, Wallops Island, Hohenpeissenberg, and Sapporo). However, the O 3 seasonal variations over these sites in Exp-flux are similar to the variations at 200 hpa (not shown). Given that at 200 hpa the O 3 concentrations (not shown) are too small in the control experiment but about the right values in Exp-flux, the model O 3 behavior at 300 hpa reflects the strong influence of stratospheric influx due to the too efficient downward transport of O 3. Over Lauder, the seasonal variation at 200 hpa (not shown) is reasonable compared with the observations, but the O 3 concentrations are overestimated at 300 hpa in both the control and Exp-flux experiments. This is also related to the too efficient downward transport caused by the SPITFIRE scheme. In the tropics, where the influence of stratospheric influx is less and the upward transport dominates, the model can reproduce the O 3 concentrations in the upper troposphere (Hilo, Natal, and Samoa), except that the smaller modeled O 3 concentrations over Natal during September November is probably related to the underestimation of the lower tropospheric biomass burning activity. [35] Figure 9 shows that the model can reasonably simulate the vertical O 3 profiles over several sonde sites for January and July. Detailed discrepancies are consistent with the above discussion. The larger model O 3 concentrations in the boundary layer over Hohenpeissenberg, Sapporo, and Wallops Island further confirm the weaker local dry depositions. Figure 10 shows the geographical O 3 distributions at 300, 500 hpa, and the surface for January and July. At 300 hpa, O 3 concentrations are relatively small in the tropics and increase toward higher latitudes with steep latitudinal gradients in the subtropics. The minimum of O 3 concentrations is located over the tropical western Pacific, the warm pool area, because of the local ascending air motion. Over the tropical Atlantic, O 3 concentrations are largest along the tropical latitude band, consistent with the tropical longitudinal variations of the inferred tropospheric O 3 by Hudson and Thompson [1998]. In the extratropics, the O 3 distribution is relatively zonally symmetric. The concentrations are larger in the summer hemisphere because of the accumulation throughout the winter to spring when the downward flux reaches its seasonal peak. At 500 hpa, O 3 is transported from both the upper troposphere and the boundary layer. As a result, the NH O 3 concentrations are higher than the SH concentrations. At the surface, the O 3 levels are high over areas with net O 3 production. In the NH during January, production of O 3 is large near the subtropics because of the relatively abundant solar irradiance, and the produced O 3 can be transported farther from its sources because of its relatively longer lifetime. During July, high concentrations are located over continents because anthropogenic emissions are high and the lifetime of O 3 is shorter Seasonal Variations of the O 3 Budget [36] Figure 11 shows the seasonal variations of tropospheric O 3 budget for the 3-year control run. The budget terms include the tropospheric O 3 tendencies due to dry deposition (solid lines), stratospheric influx (long-dashed lines), and net chemical production (short-dashed lines). The terms are integrated in the troposphere for the whole globe, the NH, and the SH. The tropopause is defined as the level with 150 ppbv O 3 concentrations for the integrations. In the NH, the stratospheric O 3 influx peaks in winter to spring and reaches its minima during summer to autumn, consistent with the temporal variations of stratospheric planetary wave activity. The net chemical production is always positive and peaks during summer, because of the high NH emissions of pollutants and the seasonal variation of solar irradiance in the NH. The dry deposition follows the variation of the surface O 3 concentrations and peaks in late summer when O 3 is accumulated through photochemical 13 of 27

14 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 8. Seasonal variations of O 3 concentrations (in ppbv) for the 3 control model years (dashed lines) and for Exp-flux (thick dashed lines; for definition, see the footnote of Table 5) compared with sonde climatology [Logan, 1999] (solid lines) at (top) 300 hpa, (middle) 500 hpa, and (bottom) 800 hpa. The location of each sonde site is shown in the title of each column. They are Edmonton (114 W, 53 N), Hohenpeissenberg (11 E, 48 N), Sapporo (141 E, 43 N), Boulder (105 W, 40 N), Wallops Island (76 W, 38 N), Kagoshima (131 E, 32 N), Hilo (155 W, 20 N), Natal (35 W, 6 S), Samoa (170 W, 14 S), and Lauder (170 E, 45 S). The vertical bars show the interannual variability (standard deviation of different years data) of the sonde climatology. 14 of 27

15 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 8. (continued) 15 of 27

16 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 8. (continued) 16 of 27

17 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 9. Vertical profiles of O 3 concentrations (in ppbv) for the 3 control model years (dashed lines) compared with sonde climatology [Logan, 1999] (solid lines) for (top) January and (bottom) July. The location of each sonde site is shown in the title of each column. Detailed latitudes and longitudes of the sites are included in the caption of Figure 8. The horizontal bars show the interannual variability of the sonde climatology. 17 of 27

18 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 9. (continued) 18 of 27

19 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 10. Geographical O 3 distributions at 300, 500 hpa and the surface (in ppbv) for January and July from the control climatologies. 19 of 27

20 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 11. Seasonal variations of the (top) global, (middle) NH, and (bottom) SH tropospheric O 3 budget for the 3 control model years. The solid lines represent the variations of dry deposition rates (in Tg/yr); the shortdashed lines represent the variations of net photochemical production (in Tg/yr); and the long-dashed lines represent the variations of stratospheric influx (in Tg/yr). production. In the SH, the stratospheric influx peaks during austral winter (May August) and has smaller temporal variability compared to the NH influx, consistent with the fact that the SH has weaker stratospheric planetary wave activity [Wong et al., 1999]. During austral summer to autumn, the net chemical production is negative when HO 2 concentrations and O 3 losses peak. The net production is positive during late austral winter to early spring (August October) because of the emissions of O 3 precursors from biomass burning. The dry deposition again follows the surface O 3 concentrations and peaks in late austral winter to early spring. The variation in the global budget is similar to that of the NH budget. [37] Table 5 illustrates the annually and globally averaged tropospheric O 3 budget for our model experiments as well as for other models from the recent literature. The globally and annually averaged OH concentrations and CO burdens from the 3 experiments are also shown. More detailed comparisons of O 3 budgets between CTMs and coupled GCMs are given by Shindell et al. [2001a, Table 5]. From these tables it can be concluded that there are still large uncertainties in tropospheric O 3 simulations. For the stratospheric influx, our control run prescribes a downward flux with a value of 600 Tg/yr, which is about in the middle of the range spanned by different models values. With this value of stratospheric influx, our model s net chemical production of tropospheric O 3 becomes 495 Tg/yr, which is at the upper end of the range spanned by different models values. The dry deposition rate becomes 1103 Tg/yr, which is also in the upper end of the range spanned by different model s values (note that our dry deposition rates are already weak over some sonde sites, see section 4.1). Doubling the production of NO by lightning results in an increase in the net chemical production (of about 18 Tg/yr or 3.6% of the control value). A detailed analysis showed that the increase in chemical production mainly occurs in the upper troposphere. The tropospheric O 3 burden increases by about 19 Tg. With an increase in the averaged OH concentration (10.6% of the control value), the CO burden decreases by about 4% of the control value. Using the prescribed stratospheric O 3, the stratospheric O 3 influx is enhanced by about 311 Tg/yr (52% of the control value). The net O 3 production decreases by about 270 Tg/yr (54% of the control value) because of the enhancement in tropospheric O 3 losses (which are proportional to the O 3 burden), and the tropospheric O 3 burden increases by about 18 Tg, almost the same amount as that due to the doubling in NO production by lightning. The averaged OH concentration is increased by 3% of the control value with a decrease of the CO burden by 2% of the control value. The comparison among different sensitivity runs shows that tropospheric oxidation capacity, represented by the globally and annually averaged OH concentration, is more sensitive to the upper tropospheric NO production by lightning than to the stratospheric O 3 influx. In Exp-flux, large O 3 changes due to the enhanced stratospheric influx occur in the extratropical upper troposphere, where, compared to the tropical region, the air is relatively dry and the solar irradiance is relatively weak. In Exp-ligtn, large O 3 changes due to reactions of the increased NO with peroxy radicals occur in the tropical upper troposphere, where production of HO x is more efficient because of the greater abundance of water vapor and the stronger solar irradiance. Increased levels of NO also shift the HO x partition toward OH because of the reaction of NO + HO 2. As a result, the abundance of tropospheric OH is sensitive to the increased abundance of NO in the tropical upper troposphere caused by greater lightning activity. 5. Radiative Forcing of O 3 Change Since the Preindustrial Period [38] The change in O 3 from the preindustrial period to the present-day is estimated by performing simulations for both present-day and preindustrial emissions. For the present-day conditions, we use the results from Exp-ligtn (with the total NO production by lightning set to be 6 TgN/yr) as it reproduces more reasonable OH and CO contents. For the preindustrial conditions, emissions from fossil fuel burning 20 of 27

21 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Table 5. Globally and Annually Averaged O 3 Budget, OH Concentrations, and CO Burden a Models/Experiments Strat. Flux Net Chem. Dry Deposition Burden [OH] CO Control Exp-ligtn b Exp-flux c Mickley et al. [1999] Lelieveld and Dentener [2000] Shindell et al. [2001a] a O 3 budget is given in Tg/yr for tendency and Tg for burden, OH concentrations are given in 10 5 molecules/cm 3, and CO burden is given in Tg. For a detailed list of all the relevant studies, please refer to Table 5 of Shindell et al. [2001a] and IPCC [2001]. Strat. Flux, stratospheric O 3 influx. Net Chem., net chemical production of O 3. b Exp-ligtn is similar to the control experiment but with the NO production by lightning increased to 3 Tg N/yr from 6 Tg N/yr. c Exp-flux is similar to the control experiment but with the stratospheric O 3 concentrations prescribed by observations. are removed, and those from biomass burning are reduced to 1/10 of the corresponding present-day values [Mickley et al., 1999; Shindell et al., 2001a]. CH 4 concentrations are set to be 0.7 ppmv for calculations of chemistry and the prescribed k coefficients for heterogeneous removal of N 2 O 5 are reduced by regionally dependent factors (0.15 for the NH extratropics, 0.24 for the tropics, and 0.7 for the SH extratropics) that are simply estimated from a model fraction of sulfate derived from natural sources [Koch et al., 1999] to account for the smaller aerosol loading. Two experiments are conducted to test the influence of lightning because of its larger effects on tropical upper tropospheric O 3 (see Figure 13) and the relatively higher sensitivity of radiative forcing to the tropical upper tropospheric O 3. The first one (Exp-P6) has the total NO production by lightning set to be 6 TgN/yr, while the second one (Exp-P3) has the Figure 12. Geographical distributions of the (top) surface O 3 concentrations (ppbv) in the preindustrial period for January and July, and the (bottom) changes in tropospheric column O 3 (in DU) during December February and June August since the preindustrial period for Exp-P6 (for definition, see the footnote of Table 6). For the lower row, the numbers next to the season labels are the global averaged column O 3 changes. 21 of 27

22 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 13. Changes in zonal mean O 3 concentrations (in percent) since the preindustrial period for (top) Exp-P6 (for definition, see the footnote of Table 6) and (bottom) Exp-P3 (for definition see the footnote of Table 6) in January and July. production set to be 3 TgN/yr, assuming that a cooler preindustrial climate may have weaker penetrative convection. For both experiments, SSTs, CO 2, N 2 O and CH 4 concentrations for radiative transfer calculations were kept the same as the present-day values while changes in O 3 concentrations below 100 hpa were allowed to influence the radiative transfer. Analyses of the results show insignificant temperature changes between the present-day and preindustrial simulations. [39] The GCCM simulated O 3 changes between the preindustrial and present-day periods were then used as inputs for off-line calculations of instantaneous radiative forcings. The background climate state for the off-line computation used the monthly mean meteorological fields in Exp-ligtn. The CCM3 s radiative transfer codes were applied on the middle day of each month and run at 0000, 0400, 0800, 1200, 1600, and 2000 GMT to account for the diurnal migration of solar irradiance. The radiative fluxes of the six time steps were averaged for the month. For the radiative fluxes in the preindustrial period, only tropospheric O 3 concentrations are perturbed by the monthly mean O 3 difference between the preindustrial and present periods, and the other climate parameters are kept unchanged. The instantaneous radiative forcings are derived by subtracting the fluxes of the perturbed state from those of the background state. In the calculation, the instantaneous radiative forcings are evaluated at the tropopause, which is defined as the level at which O 3 concentrations are 150 ppb Table 6. Annually Averaged Changes in Column O 3, CO Burden, and Globally Averaged OH Concentration in the Troposphere Since the Preindustrial Period and the Associated Instantaneous Total Radiative Forcings DO 3,DU DCO, Tg DOH, 10 5 molecules/cm 3 Radiative Forcing, Wm 2 Normalized Forcing, Wm 2 /DU Exp-P6 a Exp-P3 b a Exp-P6 assumes that the preindustrial period has the same rate of NO production from lightning as the present-day period (i.e., 6 Tg N/yr). b Exp-P3 assumes that the preindustrial period has half the rate of NO production from lightning as the present-day period (i.e., 3 Tg N/yr). 22 of 27

23 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 tions using monthly mean inputs and online calculations using instantaneous inputs for every time step in the month is small (less than 2%). We have also tested the sensitivity of the instantaneous forcings to the choice of day in the month and found that its effect was negligible. [40] The model biases of O 3 concentrations mentioned in section 4.1 can generate errors in the estimates of radiative forcings. The overestimates of O 3 concentrations in the NH middle-upper troposphere are mainly due to the too efficient downward transport of stratospheric O 3. As the stratospheric O 3 influx is the same in the preindustrial and present-day periods in our experiments, this effect can be partly cancelled out to minimize the errors in the radiative forcing. In summer, most of the overestimates are comparable to the interannual variation of O 3 concentrations and located within the bottom 1 2 km near the surface. As the contribution to radiative forcing from O 3 near the surface is relatively small, the errors generated by the overestimation should also be small. Figure 14. Geographical distributions of (top) the annual mean, (middle) December February averaged, and (bottom) June August averaged total instantaneous radiative forcings (in W m 2 ) caused by the O 3 changes since the preindustrial period for Exp-P6 (for definition, see the footnote of Table 6). The numbers in the titles are the globally averaged forcings. in the background climate state. Both perturbed O 3 concentrations in Exp-P6 and Exp-P3 were used to calculate the radiative forcings so that the uncertainty caused by the uncertain NO production by lightning could be assessed. Mickley et al. [1999] showed that the difference in the instantaneous radiative forcings between off-line calcula Perturbed O 3 Concentrations [41] Figure 12 shows the surface O 3 concentrations in Exp-P6 (upper panel) for January and July, and the associated changes in tropospheric column O 3 (lower panel) between Exp-ligtn and Exp-P6 for December January February (DJF) and June July August (JJA) seasons. Our simulations of the surface distribution of O 3 in the preindustrial period are in general similar to those shown by Mickley et al. [1999, Figure 9] and Shindell et al. [2001a, Figure 12]. Surface O 3 concentrations are larger over the wintertime oceans because of the stronger stratospheric influx and the smaller photochemical loss and dry deposition. Over continents, maxima are associated with highaltitude terrains (e.g., over Greenland and Himalaya) and the remaining biogenic emissions and biomass burning. Mickley et al. [2001] presented observations of preindustrial O 3 levels over a few sites. In all the observed data shown, the preindustrial O 3 levels are lower than 15 ppbv. Our simulated preindustrial O 3 concentrations are generally over 15 ppbv in the winter hemispheres and share the same problem as other models in overestimating the preindustrial O 3 levels. Origins of the problem are still unclear in the community, probably because of the uncertainty in the input parameters to simulate the preindustrial tropospheric chemistry [Mickley et al., 2001]. It should also be noted that calibration problems exist in the observations of the preindustrial O 3 [Kley et al., 1988; Pavelin et al., 1999]. [42] Tropospheric column O 3 is computed as the vertically integrated O 3 concentrations below the tropopause (1 DU = molecules/cm 2 ). In DJF, large changes are located near the NH subtropics because of the relatively more abundant solar irradiance. In JJA, most areas in the NH middle latitudes have changes in column O 3 over 15 DU with maxima reaching DU over the eastern United States and Mediterranean. The patterns of the changes are roughly in agreement with those given by Mickley et al. [1999], but our model has larger magnitudes of the changes. [43] Changes of the zonal mean O 3 concentrations in Exp-ligtn from Exp-P6 and Exp-P3 are shown in Figure 13 for January and July. Changes of O 3 concentrations are larger in the NH, consistent with the fact that most anthro- 23 of 27

24 WONG ET AL.: GCCM SIMULATIONS OF TROPOSPHERIC O 3 Figure 15. (top) Longwave and (bottom) shortwave components of the total instantaneous radiative forcings for the December February averages (left panels) and the June August averages (right panels) for Exp-P6 (for definition, see the footnote of Table 6). The numbers in the titles are the globally averaged forcings. pogenic emissions of precursors occur in the NH. The patterns and the magnitudes of the changes are similar to the results shown by Mickley et al. [1999], with a maximum over 120% located around N in July; however, our model shows an additional maximum in the tropical-subtropical upper troposphere due to the effects of lightning. This maximum was also seen in the absolute O 3 changes given by Berntsen et al. [1997]. In July, changes in O 3 in the subtropical upper troposphere (around 200 hpa) can be as large as 100%. If NO production by lightning is smaller in the preindustrial period, the region of O 3 changes over 80% extends over a larger area. [44] Table 6 shows the annually and globally averaged changes in column O 3, OH concentration and CO burden of Exp-ligtn from Exp-P6 and Exp-P3. When SSTs are kept unchanged in both the preindustrial and present-day simulations, two opposite factors can affect the changes of OH from the preindustrial period. The increases of O 3 and NO x enhance the OH production; on the other hand, the increases of CH 4, CO and NMVOCs enhance the OH loss. The net effect, as shown in Table 6, is that OH decreases from the preindustrial levels in our simulations. More NO production by lightning in the preindustrial period leads to a higher preindustrial OH level, corresponding to a larger decrease in OH and a larger increase in CO since the preindustrial period. The simulated decrease in the averaged OH concentrations is about 18% of the preindustrial concentration for Exp-P6, comparable to the value of 16% computed by Mickley et al. [1999] but larger than the value of 5.9% computed by Shindell et al. [2001a] Instantaneous Radiative Forcings [45] Figure 14 shows the geographical distributions of the annual mean, DJF averaged and JJA averaged total instantaneous radiative forcings (shortwave + longwave). Figure 15 shows the seasonal averaged shortwave and longwave components of the total instantaneous forcings for DJF and JJA. The annual mean total radiative forcing is over 0.5 W m 2 over most of the area between 30 S and 60 N, with maxima over 1 W m 2 located over North Africa to subtropical Atlantic and around 30 N of North America. In DJF, the total forcing is significant in latitudes lower than 60 in both hemispheres, and dominated by the longwave forcing that depends on the temperature contrast between the tropopause and the surface as well as the tropospheric O 3 changes. Shortwave forcing contributes about 22% of the 24 of 27