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Impact of ozone layer depletion II:
Atmospheric Environment 35 (2001) 1615}1625
Impact of ozone layer depletion II: changes in photodissociation rates and tropospheric composition Renata De Winter-Sorkina* Institute for Marine and Atmospheric Research Utrecht (IMAU), Utrecht University, P.O. Box 80005, 3508 TA Utrecht, The Netherlands Received 11 May 2000; received in revised form 17 August 2000; accepted 29 August 2000
Abstract Seasonal trends in the ozone pro"le determined from the Nimbus-7 TOMS total ozone record and SAGE trends are used to estimate the trends in tropospheric photodissociation rates. The resultant photodissociation trends for tropospheric ozone, CH CHO, HNO , HNO and CH O are presented. A three-dimensional chemical transport model is used to investigate the impact of changing photodissociation rates on tropospheric composition. First, the &linear trend' method is applied. The trends in OH, tropospheric ozone, carbon monoxide, methane and NO due to anthropogenic V stratospheric ozone depletion are estimated from the trends in photodissociation rates assuming a mean stratospheric ozone seasonal cycle. Total ozone trends are reduced by their 2p error estimates to exclude the in#uence that measurement errors and interannual natural variations in stratospheric ozone may have on trends in tropospheric composition. The resulting lower limits of trends in tropospheric composition due to anthropogenic ozone depletion over the time period 1978}1993 are presented at a 95% con"dence level. Second, the TOMS total ozone record is used to calculate via parameterization the temporal dependence of tropospheric photodissociation rates within the chemical transport model and to identify the resulting trends in tropospheric composition. The trends in OH, O , CO, NO and V CH determined from the TOMS total ozone record agree with the trends from the &linear trend' method within the 2p error limits. 2001 Elsevier Science Ltd. All rights reserved. Keywords: Ozone trends; UV radiation; OH radical; Tropospheric ozone; Carbon monoxide
1. Introduction Ozone levels change periodically as part of regular natural cycles such as the changing seasons, sun cycles and winds. Over the past two decades human activities have been altering the ozone balance. Human production of chlorine-containing chemicals such as chloro#uorocarbons (CFCs) and halons (also sources of Br) has added additional ozone destroying mechanisms. Stratospheric ozone depletion leads to an increase in actinic UV-B #uxes (290}320 nm) in the troposphere and to the acceleration of photodissociation rates of tropospheric ozone and some other species. The photodissociation of ozone to O(D) in the presence of water vapour yields
* Corresponding author. Fax: #31-30-2543163. E-mail address: [email protected] (R. De Winter-Sorkina).
the important hydroxyl radical OH. The OH radical can be considered as the cleansing agent of the troposphere (Levy, 1971; Thompson, 1992). Its abundance determines the lifetime of several chemical species (e.g. CO, CH ) in the troposphere prior to chemical breakdown. Tropospheric OH increases with stratospheric ozone depletion (Madronich and Granier, 1992; Fuglestvedt et al., 1994; Bekki et al., 1994). The sensitivity of OH to the ozone column is higher for lower NO levels (WMO, 1999). V Ozone in the lower troposphere tends to decrease with stratospheric ozone depletion (Bekki et al., 1994), except in very polluted regions with high NO levels, where V there is an increase of ozone in the lower troposphere during spring (Fuglestvedt et al., 1994). It seems likely that stratospheric ozone depletion has contributed to the observed changes in CH and CO trends, though it cannot explain these changes entirely (Bekki et al., 1994; Granier et al., 1996).
1352-2310/01/$ - see front matter 2001 Elsevier Science Ltd. All rights reserved. PII: S 1 3 5 2 - 2 3 1 0 ( 0 0 ) 0 0 4 3 7 - 4
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The impact of stratospheric ozone on UV-B radiation and on tropospheric chemical species such as ozone has already been observed. Spectral measurements of UV-B radiation that have been made at Toronto since 1989 indicate large upward trends linked to ozone depletion (Kerr and McElroy, 1993). Schnell et al. (1991) found a 17% reduction in surface ozone at the South Pole in December, January and February over the 1976}1990 period and suggested that the principal mechanism responsible is the increased photochemical destruction of ozone resulting from increased UV penetration. Taalas et al. (1997) found that springtime stratospheric ozone loss had a pronounced impact on the upper tropospheric ozone at Marambio, Antarctica (643S) and at SodankylaK , Finland (673N). In the 6}8 km layer, the ozone sounding records show an average ozone deviation from the 1988 to 1994 means of !12.8% in Antarctica and of !10% in the Arctic during the months with stratospheric ozone loss. BroK nnimann and Neu (1998) suggested a possible photochemical link between extremely low stratospheric ozone and record high near-surface ozone under high NO conditions on Swiss mountain sites in V late winter. Earlier studies (Bekki et al., 1994; WMO, 1999) have already used the total ozone column from TOMS measurements as input for calculations of photodissociation rates. Using the 2D models and the 3D IMAGES model they calculated the change in the global tropospheric OH and O distributions, and the change in CH and CO trends over the 1979}1993 period resulting from stratospheric ozone change. Below we will refer to this type of estimate as the &TOMS ozone' method. However, the inter-annual variations in total ozone, caused by the QBO winds, the solar cycle, meteorological conditions, volcanic eruptions and stratospheric temperatures, are partly responsible for the changes found in tropospheric composition. Additionally, the TOMS instrument suffered degradation, which could have introduced a drift in the TOMS data record of up to 1% per decade (McPeters et al., 1996). We want to "nd the minimum response of the tropospheric composition to anthropogenic stratospheric ozone depletion over this time period. We assume a linear total ozone trend and a mean annual total ozone cycle. A trend error calculation, which includes the inter-annual variations in total ozone, is performed. Below we will refer to this type of estimate as the &linear trend' method. A trend analysis was carried out on the TOMS total ozone record and an ozone pro"le depletion climatology was constructed on the basis of TOMS data and SAGE trends (De WinterSorkina, 2001). This climatology was used to estimate the trends in tropospheric photodissociation rates. Here, we report and compare the trends in tropospheric OH, O , CO, NO and CH distributions from 1979 to 1993 V estimated by both the &linear trend' and &TOMS ozone' methods using the 3D MOGUNTIA model.
2. The radiative transfer model An isotropic two-stream radiative transfer model (De Winter-Sorkina and Van der Woerd, 1993) calculates atmospheric solar photon #uxes and is used to determine the tropospheric photodissociation rates for various chemical species for clear-sky conditions. This radiative transfer model is used in a di!erence mode to investigate the impact of changing ozone on actinic #ux and photodissociation rates in the troposphere. First, the model uses monthly mean TOMS ozone column "elds for 1978/1979 (seasonal term with e!ects of QBO, solar cycle and month-to-month correlation of ozone values excluded, see De Winter-Sorkina, 2001) to estimate the monthly mean photodissociation rates for 1978/1979. Next, the ozone pro"le reduction is applied according to the climatology of monthly ozone trend pro"les (De Winter-Sorkina, 2000) to calculate the monthly mean photodissociation rates for 1992/1993. Photodissociation rates are calculated using quantum yields from DeMore et al. (1997) every 15 min during the day of the 15th of each month, and daytime or diurnal averaging is performed. Finally, the monthly trends in daytime mean photodissociation rates are calculated as a di!erence between photodissociation rate levels at the beginning and at the end of the Nimbus-7/TOMS record. Here, we assume that the e!ects of possible trends in cloud coverage, tropospheric ozone and aerosols are negligible over the period under consideration. The TOMS re#ectivity data indicate there has been no signi"cant change in zonal average aerosol plus cloudiness from 1979 to 1992 (Herman et al., 1996). This result is supported by the independent estimate of cloud e!ects based on 5 years of ERBE data (Lubin and Jensen, 1995). Possible tropospheric ozone changes are only partially visible in TOMS data (about half the change), which leads to an underestimate of the corresponding UV-#ux change (Herman et al., 1996). Trends in tropospheric ozone are highly variable (Logan, 1994) and it is di$cult to make generalizations because sonde data from only a small number of stations are suitable (WMO, 1999). Long-term trends in tropospheric ozone and sulphate aerosols have contributed only slightly to surface UV trends (Liu et al., 1991; Madronich, 1992); even less e!ect is expected on mid-tropospheric UV radiation. In our study, a simple two-stream radiative transfer model was used. To estimate possible errors in trends of photodissociation rates, our model was compared with the TUV model (Madronich, 1993) using DISORT (discrete ordinate method) (Stamnes et al., 1988) with 16 streams for radiative transfer. The model DISORT and TUV had previously been shown to give good agreement with measurements (Stamnes et al., 1991; Shetter et al., 1992; Zeng et al., 1994). There is a signi"cant di!erence in the photodissociation rates calculated by TUV with DISORT and the two-stream model even when the e!ects of
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errors made by the two-stream model compared to TUV with DISORT in calculating trends in ozone photodissociation rates (% per decade) were estimated as a function of solar zenith angle, altitude, total ozone and ozone depletion. The errors in the two-stream model are within 1% at the surface, within 2% for altitudes up to 9 km, within 3.5% for altitudes up to 15 km for solar zenith angles (sza) up to 653; within 1.5% (0 km), within 2% (9 km), within 5% (15 km) for sza of 753; within 4% (9 km) and 7% (15 km) for sza of 853; within 4% (0 km), within 8% (9 km) and larger than 10% (15 km) for sza of 893; the total ozone is from 300 to 400 DU, ozone depletion is from 2 to 16%. In our comparison we used di!erent aerosol pro"les in the two models; this could have led to an overestimation of the errors in the two-stream model. Maximum possible errors in the daily mean ozone photodissociation trends were estimated for all latitude bands, months and di!erent altitudes. The average (over all latitude bands and months) maximum possible error is 0.3% (0 km), 0.4% (9 km), 2.0% (15 km) with a standard deviation of 0.2% (0 km), 0.3% (9 km), 2.1% (15 km) and the maximum error found was 1.2% (0 km), 2.4% (9 km), 19.6% (15 km). However, the maximum errors larger than a few percent were calculated only for high latitudes winter months when the TOMS data are missing due to polar night or total ozone trends are non-signi"cant (De Winter-Sorkina, 2001, see Fig. 1). Moreover, in the chemistry-transport model (see Chapter 4) the calculation of photodissociation rates is limited to a maximum solar zenith angle of 853 (Krol and Weele, 1997).
3. Trends in photodissociation rates
Fig. 1. Trends in daytime average ozone photodissociation rate coe$cients at the surface for the time period 11/1978}04/1993 (top) together with their relative errors at 2p level (middle) and the relative di!erence in trends due to the accurate altitude dependence of ozone depletion, compared to simple scaling of ozone pro"les to "t ozone columns (bottom). Shaded areas represent the trends signi"cant at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
clouds are not included. However, when the two-stream model is checked for its accuracy in calculating trends in photodissociation rates (% per decade) for clear sky due to changing ozone, it appears to satisfy our needs. The
Considerable positive trends were found for the photodissociation rates of tropospheric ozone (De WinterSorkina, 1997) in the reaction O #hl (j(320 nm)P O #O(D), acetaldehyde (CH CHO (R)#hlP CH #CO#H), nitric acid (HNO #hlPNO #OH), peroxy nitric acid (HNO #hlPHO #NO ) and formaldehyde (CH O (R)#hlPHCO#H). Fig. 1 shows the resulting trends in daytime average ozone photodissociation rate coe$cients at the surface over the time period 11/1978}04/1993 together with the relative errors at 2p level and the relative di!erence in trends (in percentages) due to the accurate altitude dependence of ozone depletion (De Winter-Sorkina, 2001), compared to the simple scaling of ozone pro"les to "t ozone columns. Trends in the daytime average photodissociation rate of ozone as high as #18.0$4.7% per decade were found in February at the surface for zonal averages between 40 and 503N. There are also signi"cant ozone photodissociation trends of #3}#15% per decade at the surface at northern mid-latitudes in spring, summer and the "rst half of autumn, which could contribute to changes in
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tropospheric chemistry. In the Southern Hemisphere, trends in ozone photodissociation rates of #42$15% per decade at the surface occur in October between 60 and 703S. The maximum ozone photodissociation rate is found in the troposphere at about 4}5 km. The trends in ozone photodissociation rates due to stratospheric ozone depletion increase with altitude, reaching their maximum in winter}spring at about 8}12 km depending on latitude (De Winter-Sorkina, 1997). The errors in daytime average ozone photodissociation trends are calculated from the ozone trends, reduced by their errors at 2p level (for the estimation of errors in ozone trends see De Winter-Sorkina, 2001). The accurate altitude dependence of ozone depletion (as described in De Winter-Sorkina, 2001), compared to the simple scaling of ozone pro"les to the total amount of ozone, leads to mainly small changes (up to 25%) in the resultant ozone photodissociation trends as shown in Fig. 1. In general, the trends in tropospheric ozone photodissociation rates are larger than the trends in total ozone which cause them.
The trends in daytime average tropospheric photodissociation rates of species other than ozone at the surface over the time period 11/1978}04/1993 are shown in Fig. 2. The trends at the surface in January}April north of about 403N were calculated to be as high as #7}#10% per decade for CH CHO (R); #6% per decade for HNO ; #4}#6% per decade for HNO ; #4}#6% per decade for CH O (R); #2% per decade for H O , CH OOH and N O (I); #1% per decade for CH O (M). The relative errors in photodissociation rates of these species are very close to the relative errors for ozone, shown in Fig. 1. The radiation ampli"cation factors (RAF) are de"ned as the percentage increase in photodissociation rates for each percentage decrease in the ozone column (Madronich and Granier, 1992) and are computed as !ln(1#¹ )/ln(1#¹ ), where ¹ and T are the frac( ( tional trends in photodissociation rates and ozone column, respectively. The daytime mean RAFs for ozone and CH CHO (R) photodissociation at the surface over the time period 11/1978}04/1993 are shown in Fig. 3 as
Fig. 2. Trends in daytime average photodissociation rate coe$cients of CH CHO (R), HNO , HNO and CH O (R) at the surface for the time period 11/1978}04/1993. Shaded areas represent the trends signi"cant at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
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Fig. 3. The RAF for the daytime mean ozone and CH CHO (R) photodissociation at the surface over the time period 11/1978}04/1993. Shaded areas represent the signi"cance area at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
a function of latitude and month. The variations in RAF occur due to its dependence on total ozone, on the altitude of the ozone perturbation and on the solar zenith angle. The RAF for daytime mean ozone photodissociation to O(D), for example, lies between 1.3 and 2.1 (see Fig. 3).
4. Impact of trends in photodissociation rates on tropospheric composition The sensitivity of tropospheric composition to stratospheric ozone depletion is estimated in two ways. The &linear trend' method "rst assumes linear trends due to the anthropogenic in#uence of CFC emissions on total ozone, photodissociation rates and tropospheric species concentrations. Natural variations in stratospheric ozone such as the QBO and the solar cycle are excluded. The mean ozone seasonal cycle is used and the natural changes of ozone from year to year are included
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in the error estimation for the &linear trend' method. The minimum sensitivity of tropospheric composition to anthropogenic stratospheric ozone depletion is estimated by reducing the total ozone trends by their 2p errors. The use of the TOMS total ozone data to calculate photodissociation rates in the chemistry-transport model allows us to follow temporal changes in the tropospheric composition, and the assumption of linear trends in species concentrations is made at the end to determine trends. The e!ect of the photodissociation on tropospheric composition is studied by means of the 3D (103 latitude by 103 longitude) MOGUNTIA chemistry-transport model (Crutzen and Zimmermann, 1991; Dentener and Crutzen, 1993). The MOGUNTIA model is used in conjunction with monthly averaged transport (Krol, 1994), the advection-di!usion scheme is taken from Vested et al. (1992), and only background CH , CO, NO , and HO V V chemistry is considered. The calculated chemical composition of the atmosphere shows roughly the correct behaviour (Krol and Weele, 1997). The MOGUNTIA model uses the parameterization of photodissociation rates to represent the e!ects of stratospheric ozone absorption, clouds and surface re#ection (Krol and Weele, 1997). The TOMS total ozone data, the ISCCPC2 monthly averaged data on clouds, on snow/ice cover (Rossow and Schi!er, 1991) and the surface albedo data set (based on land-use data (Olson and Watts, 1982) and UV-A, UV-B average albedos (Feister and Grewe, 1995)) are used to calculate the photodissociation rates by means of the parameterization scheme. To study the sensitivity of tropospheric composition to changes in photodissociation rates, the MOGUNTIA model was run for a period of 15 years from 1978 to 1993 for three di!erent cases, the emissions being kept constant. First, the &no trend' MOGUNTIA model run was performed. The photodissociation rates were calculated via parameterization using the TOMS total ozone and ISCCP cloud data for one year from the beginning of the TOMS data record. Second, the &linear trend' MOGUNTIA model run was made. To the previously described photodissociation rates, calculated using 1 year of TOMS data, we added the linear trends in photodissociation rates of di!erent species due to stratospheric ozone depletion (see Chapter 3). These trends were calculated with the radiative transfer model every 2 h (MOGUNTIA time step) for the 15th of each month for latitude bands of 103. Third, the &reduced linear trend' MOGUNTIA model run was made. The aim was to determine the lower limits of tropospheric composition sensitivity to linear trends in total ozone due to the anthropogenic in#uence of CFC emissions. The stratospheric ozone trends are reduced by their 2p errors (De Winter-Sorkina, 2001) and the reduced trends in photodissociation rates are calculated. The di!erence between this run and the previous MOGUNTIA model run lies in the use of these reduced photodissociation trends. Fourth, the
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Fig. 4. Trends in OH, O , CO and NO in the boundary layer over the time period 1979}1992 due to trends in photodissociation rates. V Shaded areas represent the signi"cance area at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
&TOMS ozone' MOGUNTIA model run was performed. The whole 14 year record of TOMS total ozone data was used to calculate photodissociation rates via parameterization. The changes in OH, O , CO, NO and V CH in time were calculated relative to the &no trend' run and modelled by the seasonal cycle, the linear trend and the noise term in a linear regression "t. The trends in OH, O , CO and NO in the boundary V layer over the time period 1979}1992 due to the trends in photodissociation rates are shown in Fig. 4. Here &linear trend' and &no trend' MOGUNTIA model runs were used. Minimum trends in OH, O , CO and NO occur V red in the boundary layer over the time period 1979}1992 due to the trends in photodissociation rates reduced by their 2p errors are shown in Fig. 5. Here &reduced linear trend' and &no trend' MOGUNTIA model runs described above were used. Trends in OH, O , CO and NO in the boundary 6 layer over the time period 1979}1992 due to the trends in photodissociation rates calculated from the TOMS ozone dataset via parameterization are shown in Fig. 6. Here &TOMS ozone' and &no trend' MOGUNTIA model
runs were used. The signi"cance area at the 2p error level is shaded and was calculated by linear regression "t. The errors in trends describe the ability to represent the temporal dependence of species concentrations by seasonal term and the linear trend. In the case of trends determined by the &linear trend' method these errors are negligible due to the use of the mean ozone seasonal cycle. Note that errors in trends from the &TOMS ozone' model run do not include the long-term drift in TOMS data. Positive trends in tropospheric OH concentration (due to stratospheric ozone depletion) in the boundary layer of northern mid-latitudes of about #6}#8% per decade (#5}#7% per decade from the &TOMS ozone' run) in February-April were calculated. The OH trend at 453N reaches its maximum of #9.3$2.6% per decade in February (#7.3$3.1% per decade in March from the &TOMS ozone' run). Note that the in#uence of this trend on the global oxidation capacity is small due to low OH. The lower limits of OH trends (when the 2p errors are subtracted from the stratospheric ozone trends, see Fig. 5) of about #4}#6 % per decade were calculated in
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Fig. 5. Minimum trends in OH, O , CO and NO in the boundary layer over the time period 1979}1992 due to trends in V photodissociation rates reduced by their 2p errors.
February}April in the boundary layer of northern midlatitudes. The OH trends estimated from the &TOMS ozone' MOGUNTIA run are somewhat lower, but within the error limits of the OH trends determined from the &linear trend' run. The annual OH trend in the boundary layer at 453N is #4.3$2.8% per decade from the &linear trend' run and #4.0$1.0% per decade from the &TOMS ozone' run. A global trend in tropospheric OH concentration due to stratospheric ozone depletion amounting to #2.4$1.3% per decade (#2.9$0.6% per decade from the &TOMS ozone' run) was calculated. This can explain part of the global linear OH trend estimated from the methylchloroform measurements to be about #4.6$0.6% per decade (Krol et al., 1998). The temporal dependence of the change in global OH calculated from the &TOMS ozone' run agrees well with the ones presented by Bekki et al. (1994) and WMO (1999). The e!ect on tropospheric ozone concentration appears to be moderate. The negative trends in tropospheric ozone concentration (due to stratospheric ozone depletion) in the boundary layer at 553N reach their maximum of !2.2$1.2% per decade (!1.8$1.0% per
decade from the &TOMS ozone' run) in June. The lower limits of tropospheric ozone trends (when the 2p errors are subtracted from the stratospheric ozone trends) of about !0.5}!1% per decade were calculated in April}July for the boundary layer at 553N. The tropospheric ozone trends estimated from the &TOMS ozone' MOGUNTIA run are within the error limits of the tropospheric O trends determined from the &linear trend' run. The annual tropospheric ozone trend in the boundary layer at 553N is !1.0$0.7% per decade from the &linear trend' run and !0.8$0.3% per decade from the &TOMS ozone' run. The global trend in tropospheric O concentration due to stratospheric ozone depletion of !0.8$0.7% per decade (!0.7$0.1% per decade from the &TOMS ozone' run) was calculated at the 2p uncertainty level. Surface measurements of CO and CH from global networks have revealed a decline in growth rates during the last decade (WMO, 1999). The rate of OH increase, which is necessary to explain the change in the CO trend at northern mid-latitudes between 1980 and 1995, is estimated to be 6$3% per decade (Yurganov et al.,
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Fig. 6. Trends in OH, O , CO and NO in the boundary layer over the time period 1979}1992 due to trends in photodissociation rates V calculated from the TOMS ozone dataset via parameterization. Shaded areas represent the signi"cance area at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
1999). The negative trends in tropospheric CO concentration (due to stratospheric ozone depletion) in the boundary layer at 453N reach their maximum of !2.2$1.1% per decade in May (!1.9$1.1% per decade in June from the &TOMS ozone' run). The lower limits of CO trends (when the 2p errors are subtracted from the stratospheric ozone trends) of about !0.5 to !1% per decade were calculated in February} June for the boundary layer at 453N. The CO trends estimated from the &TOMS ozone' MOGUNTIA run are within the error limits of the CO trends determined from the &linear trend' run. The annual CO trend in the boundary layer at 453N is !1.3$0.9% per decade from the &linear trend' run and !1.2$0.2% per decade from the &TOMS ozone' run. The global trend in tropospheric CO concentration due to stratospheric ozone depletion of !1.8$1.3% per decade (!2.0$0.2% per decade from the &TOMS ozone' run) was calculated. We might expect variation in the spatial NO chemV istry (Fuglestvedt et al., 1994) since increased OH levels will increase the transfer of NO to HNO , which is the main loss mechanism for NO . The negative trends in V
tropospheric NO concentration (due to stratospheric V ozone depletion) in the boundary layer at 553N reach their maximum of !1.7$0.4% per decade in April (!1.4$0.8% per decade in May from the &TOMS ozone' run). The lower limits of NO trends (when the V 2p errors are subtracted from the stratospheric ozone trends) of about !1% per decade were calculated in February}May for the boundary layer at 453N. The NO V trends estimated from the &TOMS ozone' MOGUNTIA run are within the error limits of the NO trends deterV mined from the &linear trend' run. The annual NO trend V in the boundary layer at 453N is !0.8$0.5% per decade from the &linear trend' run and !0.8$0.3% per decade from the &TOMS ozone' run. Urban measurements of NO in the United States show a downward trend of about !0.5% per year in the 1980s (Logan, 1994). Trend analyses of surface NO measurements in North-West V Europe over the period 1980}1994 (De Paus and Roemer, 1997) shows that concentrations of NO were V decreasing with 0.1}4.3% per year, except for UK, mainly as a result of decreasing emissions. No signi"cant global trend in tropospheric NO concentration due to V
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stratospheric ozone depletion was found at a 95% con"dence level in our study, as a result of NO trends V decreasing with altitude. Due to increases in OH, the global level of CH is reduced. Because of the long lifetime of CH , the maximum e!ect on methane is delayed by a few years (Fuglestvedt et al., 1994). The negative trends in tropospheric CH concentration (due to stratospheric ozone depletion) in the boundary layer reach their maxima of !0.8$0.7% per decade at 753N in July and !0.9$0.8% per decade at 653S in December (!1.2$0.2% per decade at 753N in August and !1.3$0.2% per decade at 753S in December from the &TOMS ozone' run). The CH trends estimated from the &TOMS ozone' MO GUNTIA run are within the error limits of the CH trends determined from the &linear trend' run. The lower limits of CH trends (when the 2p errors are subtracted from the ozone trends) of about !0.08 to 0.1% per decade were calculated for the boundary layer at 453N. The annual CH trend in the boundary layer at 453N is !0.8$0.7% per decade from the &linear trend' run and !1.15$0.05% per decade from the &TOMS ozone' run. The moderate response of methane is explained by the lack of correlation between the changes in OH (largest at higher latitudes in spring) and the rate of methane oxidation, which increases with temperature (largest at lower latitudes) (WMO, 1999). The global CH trend (at 2p error level) of !0.8$0.7% per decade (!1.2$0.1% per decade from the &TOMS ozone' run) was found over the 1979}1992 time period.
5. Discussion and conclusions An ozone depletion climatology based on the TOMS/ Nimbus-7 (11/1978}04/1993) version 7 total ozone data and SAGE ozone trends together with ozonesonde data was used to estimate monthly trends in the daytime average photodissociation rate coe$cients of tropospheric species and their daily cycle due to stratospheric ozone depletion over 103 latitude bands. The minimum response of the tropospheric composition to anthropogenic stratospheric ozone depletion over this time period is estimated. The trends in tropospheric OH, O , CO, NO and CH distributions are estimated by the &linear V trend' and the &TOMS ozone' methods using the 3D MOGUNTIA model. The trends estimated from the &TOMS ozone' method are within the error limits of the trends determined from the &linear trend' method. This means that the &linear trend' approach can be used for estimating future scenarios when the yearly ozone variations are unknown. The error analyses of stratospheric ozone depletion and its impact on tropospheric trends shows that signi"cant changes have occurred in tropospheric densities at mid- and high latitudes of both hemispheres over the
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period 1979}1992. Global trends in tropospheric concentrations due to stratospheric ozone depletion are calculated by MOGUNTIA model of #2.4$1.3% per decade for OH, !1.8$1.3% per decade for CO and !0.8$0.7% per decade for O and CH when emis sions are kept constant. Thus, a small decrease in global tropospheric ozone was estimated due to stratospheric ozone reduction. However, the e!ect of stratospheric ozone depletion on tropospheric trends is small or comparable to other factors, especially the anthropogenic emission changes. Response of global OH concentrations to stratospheric ozone loss was found to be equal to a possible 10% NO emission increase and 85% of this V response was found to be equal to an estimated 6.5% CO decrease or a possible 10% tropical H O increase (Krol et al., 1998). The estimated 11% CH increase leads to a decrease of global OH of 55% of its positive response to stratospheric ozone depletion (Krol et al., 1998). However, it should be emphasized that the changes in the global atmospheric composition since 1978 are to a large extent unknown. Granier et al. (1996) calculated CO concentration responses to combined stratospheric ozone changes over period 1989/1990}1993, tropospheric temperature changes, a 10% decrease of industrial source of CO and a 25% reduction in biomass-burning (CO, NMHC, CH and NO emissions). Under this V scenario, the changes calculated for the 1990}1993 period are consistent with the changes observed at most CO measuring stations. However, total ozone and temperature change could be responsible for no more than 30% of the total observed change in CO (Granier et al., 1996). Trends in tropospheric ozone in the context of trends in emissions of NO were discussed by Logan (1994). V Reported trends in tropospheric ozone are highly variable, sparse and sometimes ambiguous. Thus, the impact of stratospheric ozone depletion on tropospheric ozone trends can be shadowed by the impact of emission trends. However, in regions were emission changes are playing a minor role, e!ects due to stratospheric ozone reduction can be indicated. For example, three Canadian ozonesonde stations show a long-term decrease in ozone since 1980, which is signi"cant only in winter and autumn and may re#ect dynamical variability and lower amounts of stratospheric ozone (Logan, 1994). Tropospheric ozone changes due to stratospheric ozone reduction are very sensitive to local NO concentrations (Liu V and Trainer, 1988). In general, tropospheric ozone trends due to stratospheric ozone depletion are too small compared to the current uncertainty of the long-term ozonesonde and surface ozone observations to be measured.
Acknowledgements Renata de Winter-Sorkina was supported by the European Union RIFTOZ project and by the Space Research
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Organisation of the Netherlands (SRON). The author thanks Maarten Krol for performing MOGUNTIA model runs, for helpful discussions and for comments on the manuscript, and is grateful to Peter Builtjes for his interest and support.
References Bekki, S., Law, K.S., Pyle, J.A., 1994. E!ect of ozone depletion on atmospheric CH and CO concentrations. Nature 371, 595}597. BroK nnimann, S., Neu, U., 1998. A possible photochemical link between stratospheric and near-surface ozone on Swiss mountain sites in late winter. Journal of Atmospheric Chemistry 31, 299}319. Crutzen, P.J., Zimmermann, P.H., 1991. The changing photochemistry of the troposphere. Tellus 43(AB), 136}151. DeMore, W.B., Sanders, S.P., Golden, D.M., Hampson, R.F., Kurylo, M.J., Howard, C.J., Ravishankara, A.R., Kolb, C.E., Molina, M.J., 1997. Chemical kinetics and photochemical data for use in stratospheric modeling. JPL, evaluation number 12. California Institute of Technology, Pasadena, CA. Dentener, F.J., Crutzen, P.J., 1993. Reaction of N O on tropo spheric aerosols: impact on the global distributions NO , V O , and OH. Journal of Geophysical Research 98, 7149. De Paus, T.A., Roemer M.G.M., 1997. Trends of NO in NorthV west Europe. TNO Report, TNO-MEP}R 97/136, Apeldoorn, The Netherlands. De Winter-Sorkina, R., 1997. Impact of ozone layer depletion on tropospheric ozone photodissociation rates. In: Walcek, C.E. (Ed.), American Meteorological Society, Proceedings of the 3rd Conference on Atmospheric Chemistry, 2}7 February 1997, Long Beach, CA, pp. 118}127. De Winter-Sorkina, R., 2001. Impact of ozone layer depletion I: the ozone depletion climatology. Atmospheric Environment 35, 1609}1614. De Winter-Sorkina, R., Van der Woerd, H.J., 1993. Analysis of the potential of tropospheric trace gas observations from satellites. RIVM Report No. 722201001, Bilthoven, The Netherlands. Feister, U., Grewe, R., 1995. Spectral albedo measurements in the UV and visible region over di!erent types of surfaces. Photochemistry and Photobiology 62, 736}744. Fuglestvedt, J.S., Jonson, J.E., Isaksen, I.S.A., 1994. E!ects of reduction in stratospheric ozone on tropospheric chemistry through changes in photolysis rates. Tellus 46B, 172}192. Granier, C., Muller, J.F., Madronich, S., Brasseur, G.P., 1996. Possible causes for the 1990}1993 decrease in the global tropospheric CO abundances: a threedimensional sensitivity study. Atmospheric Environment 30, 1673}1682. Herman, J.R., Bhartia, P.K., Ziemke, J., Ahmad, Z., Larko, D., 1996. UV-B increases (1979}1992) from decreases in total ozone. Geophysical Research Letters 23, 2117}2120. Kerr, J.B., McElroy, C.T., 1993. Evidence for large upward trends of ultraviolet-B radiation linked to ozone depletion. Science 262, 1032}1034.
Krol, M.C., 1994. Treatment of transport in MOGUNTIA. In: Gryning, S.-E., Millan, M.M. (Eds.), Air Pollution Modelling and its Application, Vol. 18. Plenum Press, New York, pp. 623}624. Krol, M.C., Van Weele, M., 1997. Implications of variations in photodissociation rates for global tropospheric chemistry. Atmospheric Environment 31, 1257}1273. Krol, M., Van Leeuwen, P.J., Lelieveld, J., 1998. Global OH trend inferred from methylchloroform measurements. Journal of Geophysical Research 103, 10697}10711. Levy Jr., H., 1971. Normal atmosphere: large radical and formaldehyde concentrations predicted. Science 173, 141}143. Liu, S.C., McKeen, S.A., Madronich, S., 1991. E!ect of anthropogenic aerosols on biologically active ultraviolet radiation. Geophysical Research Letters 18, 2265}2268. Liu, S.C., Trainer, M., 1988. Responses of tropospheric ozone and odd hydrogen radicals to column ozone changes. Journal of Atmospheric Chemistry 6, 221}233. Logan, J.A., 1994. Trends in the vertical distribution of ozone: an analysis of ozonesonde data. Journal of Geophysical Research 99, 25553}25585. Lubin, D., Jensen, E.H., 1995. E!ects of clouds and stratospheric ozone depletion on ultraviolet radiation changes. Nature 377, 710}713. Madronich, S., 1992. Implications of recent total atmospheric ozone measurements for biologically active ultraviolet radiation reaching the Earth's surface. Geophysical Research Letters 19, 37}40. Madronich, S., 1993. UV radiation in the natural and perturbed atmosphere. In: Tevini, M. (Ed.), Environmental E!ects of UV (Ultraviolet) Radiation. Lewis Publisher, Boca Raton, pp. 17}69. Madronich, S., Granier, C., 1992. Impact of recent total ozone changes on tropospheric ozone photodissociation, hydroxyl radicals, and methane trends. Geophysical Research Letters 19, 465}467. McPeters, R.D., Bhartia, P.K., Krueger, A.J., Herman, J.R., Schlesinger, C.G., Wellemeyer, C.G., Seftor, C.J., Jaross, G., Taylor, S.L., Swissler, T., Torres, O., Labow, G., Byerly, W., Cebula, R.P., 1996. Nimbus 7 Total Ozone Mapping Spectrometer (TOMS) data products user's guide. NASA Reference Publication 1384. Olson, J., Watts, J.A., 1982. Major world ecosystem complexes. In: Clark, W.C. (Ed.), Carbon Dioxide Review. Oxford University Press, Oxford, pp. 388}399. Rossow, W.B., Schi!er, R.A., 1991. ISCCP Cloud Data Products. Bulletin of the American Meteorological Society 72, 2}20. Shetter, R.E., McDaniel, A.H., Cantrell, C.A., Madronich, S., Calvert, J.G., 1992. Actinometer and Eppley radiometer measurements of the NO photolysis rate coe$ cient during the Mauna Loa Observatory Photochemistry Experiment. Journal of Geophysical Research 97, 10349}10359. Schnell, R.C., Liu, S.C., Oltmans, S.J., Stone, R.S., Hofmann, D.J., Dutton, E.G., Deshler, T., Sturges, W.T., Harder, J.W., Sewell, S.D., Trainer, M., Harris, J.M., 1991. Decrease of summer tropospheric ozone concentrations in Antarctica. Nature 351, 726}729.
R. De Winter-Sorkina / Atmospheric Environment 35 (2001) 1615}1625 Stamnes, K., Slusser, J., Bowen, M., 1991. Derivation of total ozone column abundance from spectral irradiance measurements. Applied Optics 30, 4418}4426. Stamnes, K., Tsay, S., Wiscombe, W., Jayaweera, K., 1988. Numerically stable algorithm for discrete-ordinate-method radiative transfer in multiple scattering and emitting layered media. Applied Optics 27, 2502}2509. Taalas, P., Damski, J., KyroK , E., Ginzburg, M., Talamoni, G., 1997. E!ect of stratospheric ozone variations on UV radiation and on tropospheric ozone at high latitudes. Journal of Geophysical Research 102, 1533}1539. Thompson, A.M., 1992. The oxidizing capacity of the earth's atmosphere: probable past and future changes. Science 256, 1157}1165.
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Vested, H.J., Justesen, P., Ekebjaerg, L., 1992. Advection}dispersion modelling in three dimensions. Applied Mathematical Modelling 16, 506}519. WMO, 1999. Scienti"c assessment of ozone depletion: 1998, Global ozone research and monitoring project, World Meteorological Organization, Report No. 44, Geneva. Yurganov, L.N., Grechko, E.I., Dzhola, A.V., 1999. Zvenigorod carbon monoxide total column time series: 27 yr of measurements. Chemosphere: Global Change Science 1, 127}136. Zeng, J., McKenzie, R., Stamnes, K., Wineland, M., Rosen, J., 1994. Measured UV spectra compared with discrete ordinate method simulations. Journal of Geophysical Research 99, 23019}23030.
Impact of ozone layer depletion II: changes in photodissociation rates and tropospheric composition Renata De Winter-Sorkina* Institute for Marine and Atmospheric Research Utrecht (IMAU), Utrecht University, P.O. Box 80005, 3508 TA Utrecht, The Netherlands Received 11 May 2000; received in revised form 17 August 2000; accepted 29 August 2000
Abstract Seasonal trends in the ozone pro"le determined from the Nimbus-7 TOMS total ozone record and SAGE trends are used to estimate the trends in tropospheric photodissociation rates. The resultant photodissociation trends for tropospheric ozone, CH CHO, HNO , HNO and CH O are presented. A three-dimensional chemical transport model is used to investigate the impact of changing photodissociation rates on tropospheric composition. First, the &linear trend' method is applied. The trends in OH, tropospheric ozone, carbon monoxide, methane and NO due to anthropogenic V stratospheric ozone depletion are estimated from the trends in photodissociation rates assuming a mean stratospheric ozone seasonal cycle. Total ozone trends are reduced by their 2p error estimates to exclude the in#uence that measurement errors and interannual natural variations in stratospheric ozone may have on trends in tropospheric composition. The resulting lower limits of trends in tropospheric composition due to anthropogenic ozone depletion over the time period 1978}1993 are presented at a 95% con"dence level. Second, the TOMS total ozone record is used to calculate via parameterization the temporal dependence of tropospheric photodissociation rates within the chemical transport model and to identify the resulting trends in tropospheric composition. The trends in OH, O , CO, NO and V CH determined from the TOMS total ozone record agree with the trends from the &linear trend' method within the 2p error limits. 2001 Elsevier Science Ltd. All rights reserved. Keywords: Ozone trends; UV radiation; OH radical; Tropospheric ozone; Carbon monoxide
1. Introduction Ozone levels change periodically as part of regular natural cycles such as the changing seasons, sun cycles and winds. Over the past two decades human activities have been altering the ozone balance. Human production of chlorine-containing chemicals such as chloro#uorocarbons (CFCs) and halons (also sources of Br) has added additional ozone destroying mechanisms. Stratospheric ozone depletion leads to an increase in actinic UV-B #uxes (290}320 nm) in the troposphere and to the acceleration of photodissociation rates of tropospheric ozone and some other species. The photodissociation of ozone to O(D) in the presence of water vapour yields
* Corresponding author. Fax: #31-30-2543163. E-mail address: [email protected] (R. De Winter-Sorkina).
the important hydroxyl radical OH. The OH radical can be considered as the cleansing agent of the troposphere (Levy, 1971; Thompson, 1992). Its abundance determines the lifetime of several chemical species (e.g. CO, CH ) in the troposphere prior to chemical breakdown. Tropospheric OH increases with stratospheric ozone depletion (Madronich and Granier, 1992; Fuglestvedt et al., 1994; Bekki et al., 1994). The sensitivity of OH to the ozone column is higher for lower NO levels (WMO, 1999). V Ozone in the lower troposphere tends to decrease with stratospheric ozone depletion (Bekki et al., 1994), except in very polluted regions with high NO levels, where V there is an increase of ozone in the lower troposphere during spring (Fuglestvedt et al., 1994). It seems likely that stratospheric ozone depletion has contributed to the observed changes in CH and CO trends, though it cannot explain these changes entirely (Bekki et al., 1994; Granier et al., 1996).
1352-2310/01/$ - see front matter 2001 Elsevier Science Ltd. All rights reserved. PII: S 1 3 5 2 - 2 3 1 0 ( 0 0 ) 0 0 4 3 7 - 4
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The impact of stratospheric ozone on UV-B radiation and on tropospheric chemical species such as ozone has already been observed. Spectral measurements of UV-B radiation that have been made at Toronto since 1989 indicate large upward trends linked to ozone depletion (Kerr and McElroy, 1993). Schnell et al. (1991) found a 17% reduction in surface ozone at the South Pole in December, January and February over the 1976}1990 period and suggested that the principal mechanism responsible is the increased photochemical destruction of ozone resulting from increased UV penetration. Taalas et al. (1997) found that springtime stratospheric ozone loss had a pronounced impact on the upper tropospheric ozone at Marambio, Antarctica (643S) and at SodankylaK , Finland (673N). In the 6}8 km layer, the ozone sounding records show an average ozone deviation from the 1988 to 1994 means of !12.8% in Antarctica and of !10% in the Arctic during the months with stratospheric ozone loss. BroK nnimann and Neu (1998) suggested a possible photochemical link between extremely low stratospheric ozone and record high near-surface ozone under high NO conditions on Swiss mountain sites in V late winter. Earlier studies (Bekki et al., 1994; WMO, 1999) have already used the total ozone column from TOMS measurements as input for calculations of photodissociation rates. Using the 2D models and the 3D IMAGES model they calculated the change in the global tropospheric OH and O distributions, and the change in CH and CO trends over the 1979}1993 period resulting from stratospheric ozone change. Below we will refer to this type of estimate as the &TOMS ozone' method. However, the inter-annual variations in total ozone, caused by the QBO winds, the solar cycle, meteorological conditions, volcanic eruptions and stratospheric temperatures, are partly responsible for the changes found in tropospheric composition. Additionally, the TOMS instrument suffered degradation, which could have introduced a drift in the TOMS data record of up to 1% per decade (McPeters et al., 1996). We want to "nd the minimum response of the tropospheric composition to anthropogenic stratospheric ozone depletion over this time period. We assume a linear total ozone trend and a mean annual total ozone cycle. A trend error calculation, which includes the inter-annual variations in total ozone, is performed. Below we will refer to this type of estimate as the &linear trend' method. A trend analysis was carried out on the TOMS total ozone record and an ozone pro"le depletion climatology was constructed on the basis of TOMS data and SAGE trends (De WinterSorkina, 2001). This climatology was used to estimate the trends in tropospheric photodissociation rates. Here, we report and compare the trends in tropospheric OH, O , CO, NO and CH distributions from 1979 to 1993 V estimated by both the &linear trend' and &TOMS ozone' methods using the 3D MOGUNTIA model.
2. The radiative transfer model An isotropic two-stream radiative transfer model (De Winter-Sorkina and Van der Woerd, 1993) calculates atmospheric solar photon #uxes and is used to determine the tropospheric photodissociation rates for various chemical species for clear-sky conditions. This radiative transfer model is used in a di!erence mode to investigate the impact of changing ozone on actinic #ux and photodissociation rates in the troposphere. First, the model uses monthly mean TOMS ozone column "elds for 1978/1979 (seasonal term with e!ects of QBO, solar cycle and month-to-month correlation of ozone values excluded, see De Winter-Sorkina, 2001) to estimate the monthly mean photodissociation rates for 1978/1979. Next, the ozone pro"le reduction is applied according to the climatology of monthly ozone trend pro"les (De Winter-Sorkina, 2000) to calculate the monthly mean photodissociation rates for 1992/1993. Photodissociation rates are calculated using quantum yields from DeMore et al. (1997) every 15 min during the day of the 15th of each month, and daytime or diurnal averaging is performed. Finally, the monthly trends in daytime mean photodissociation rates are calculated as a di!erence between photodissociation rate levels at the beginning and at the end of the Nimbus-7/TOMS record. Here, we assume that the e!ects of possible trends in cloud coverage, tropospheric ozone and aerosols are negligible over the period under consideration. The TOMS re#ectivity data indicate there has been no signi"cant change in zonal average aerosol plus cloudiness from 1979 to 1992 (Herman et al., 1996). This result is supported by the independent estimate of cloud e!ects based on 5 years of ERBE data (Lubin and Jensen, 1995). Possible tropospheric ozone changes are only partially visible in TOMS data (about half the change), which leads to an underestimate of the corresponding UV-#ux change (Herman et al., 1996). Trends in tropospheric ozone are highly variable (Logan, 1994) and it is di$cult to make generalizations because sonde data from only a small number of stations are suitable (WMO, 1999). Long-term trends in tropospheric ozone and sulphate aerosols have contributed only slightly to surface UV trends (Liu et al., 1991; Madronich, 1992); even less e!ect is expected on mid-tropospheric UV radiation. In our study, a simple two-stream radiative transfer model was used. To estimate possible errors in trends of photodissociation rates, our model was compared with the TUV model (Madronich, 1993) using DISORT (discrete ordinate method) (Stamnes et al., 1988) with 16 streams for radiative transfer. The model DISORT and TUV had previously been shown to give good agreement with measurements (Stamnes et al., 1991; Shetter et al., 1992; Zeng et al., 1994). There is a signi"cant di!erence in the photodissociation rates calculated by TUV with DISORT and the two-stream model even when the e!ects of
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errors made by the two-stream model compared to TUV with DISORT in calculating trends in ozone photodissociation rates (% per decade) were estimated as a function of solar zenith angle, altitude, total ozone and ozone depletion. The errors in the two-stream model are within 1% at the surface, within 2% for altitudes up to 9 km, within 3.5% for altitudes up to 15 km for solar zenith angles (sza) up to 653; within 1.5% (0 km), within 2% (9 km), within 5% (15 km) for sza of 753; within 4% (9 km) and 7% (15 km) for sza of 853; within 4% (0 km), within 8% (9 km) and larger than 10% (15 km) for sza of 893; the total ozone is from 300 to 400 DU, ozone depletion is from 2 to 16%. In our comparison we used di!erent aerosol pro"les in the two models; this could have led to an overestimation of the errors in the two-stream model. Maximum possible errors in the daily mean ozone photodissociation trends were estimated for all latitude bands, months and di!erent altitudes. The average (over all latitude bands and months) maximum possible error is 0.3% (0 km), 0.4% (9 km), 2.0% (15 km) with a standard deviation of 0.2% (0 km), 0.3% (9 km), 2.1% (15 km) and the maximum error found was 1.2% (0 km), 2.4% (9 km), 19.6% (15 km). However, the maximum errors larger than a few percent were calculated only for high latitudes winter months when the TOMS data are missing due to polar night or total ozone trends are non-signi"cant (De Winter-Sorkina, 2001, see Fig. 1). Moreover, in the chemistry-transport model (see Chapter 4) the calculation of photodissociation rates is limited to a maximum solar zenith angle of 853 (Krol and Weele, 1997).
3. Trends in photodissociation rates
Fig. 1. Trends in daytime average ozone photodissociation rate coe$cients at the surface for the time period 11/1978}04/1993 (top) together with their relative errors at 2p level (middle) and the relative di!erence in trends due to the accurate altitude dependence of ozone depletion, compared to simple scaling of ozone pro"les to "t ozone columns (bottom). Shaded areas represent the trends signi"cant at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
clouds are not included. However, when the two-stream model is checked for its accuracy in calculating trends in photodissociation rates (% per decade) for clear sky due to changing ozone, it appears to satisfy our needs. The
Considerable positive trends were found for the photodissociation rates of tropospheric ozone (De WinterSorkina, 1997) in the reaction O #hl (j(320 nm)P O #O(D), acetaldehyde (CH CHO (R)#hlP CH #CO#H), nitric acid (HNO #hlPNO #OH), peroxy nitric acid (HNO #hlPHO #NO ) and formaldehyde (CH O (R)#hlPHCO#H). Fig. 1 shows the resulting trends in daytime average ozone photodissociation rate coe$cients at the surface over the time period 11/1978}04/1993 together with the relative errors at 2p level and the relative di!erence in trends (in percentages) due to the accurate altitude dependence of ozone depletion (De Winter-Sorkina, 2001), compared to the simple scaling of ozone pro"les to "t ozone columns. Trends in the daytime average photodissociation rate of ozone as high as #18.0$4.7% per decade were found in February at the surface for zonal averages between 40 and 503N. There are also signi"cant ozone photodissociation trends of #3}#15% per decade at the surface at northern mid-latitudes in spring, summer and the "rst half of autumn, which could contribute to changes in
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tropospheric chemistry. In the Southern Hemisphere, trends in ozone photodissociation rates of #42$15% per decade at the surface occur in October between 60 and 703S. The maximum ozone photodissociation rate is found in the troposphere at about 4}5 km. The trends in ozone photodissociation rates due to stratospheric ozone depletion increase with altitude, reaching their maximum in winter}spring at about 8}12 km depending on latitude (De Winter-Sorkina, 1997). The errors in daytime average ozone photodissociation trends are calculated from the ozone trends, reduced by their errors at 2p level (for the estimation of errors in ozone trends see De Winter-Sorkina, 2001). The accurate altitude dependence of ozone depletion (as described in De Winter-Sorkina, 2001), compared to the simple scaling of ozone pro"les to the total amount of ozone, leads to mainly small changes (up to 25%) in the resultant ozone photodissociation trends as shown in Fig. 1. In general, the trends in tropospheric ozone photodissociation rates are larger than the trends in total ozone which cause them.
The trends in daytime average tropospheric photodissociation rates of species other than ozone at the surface over the time period 11/1978}04/1993 are shown in Fig. 2. The trends at the surface in January}April north of about 403N were calculated to be as high as #7}#10% per decade for CH CHO (R); #6% per decade for HNO ; #4}#6% per decade for HNO ; #4}#6% per decade for CH O (R); #2% per decade for H O , CH OOH and N O (I); #1% per decade for CH O (M). The relative errors in photodissociation rates of these species are very close to the relative errors for ozone, shown in Fig. 1. The radiation ampli"cation factors (RAF) are de"ned as the percentage increase in photodissociation rates for each percentage decrease in the ozone column (Madronich and Granier, 1992) and are computed as !ln(1#¹ )/ln(1#¹ ), where ¹ and T are the frac( ( tional trends in photodissociation rates and ozone column, respectively. The daytime mean RAFs for ozone and CH CHO (R) photodissociation at the surface over the time period 11/1978}04/1993 are shown in Fig. 3 as
Fig. 2. Trends in daytime average photodissociation rate coe$cients of CH CHO (R), HNO , HNO and CH O (R) at the surface for the time period 11/1978}04/1993. Shaded areas represent the trends signi"cant at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
R. De Winter-Sorkina / Atmospheric Environment 35 (2001) 1615}1625
Fig. 3. The RAF for the daytime mean ozone and CH CHO (R) photodissociation at the surface over the time period 11/1978}04/1993. Shaded areas represent the signi"cance area at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
a function of latitude and month. The variations in RAF occur due to its dependence on total ozone, on the altitude of the ozone perturbation and on the solar zenith angle. The RAF for daytime mean ozone photodissociation to O(D), for example, lies between 1.3 and 2.1 (see Fig. 3).
4. Impact of trends in photodissociation rates on tropospheric composition The sensitivity of tropospheric composition to stratospheric ozone depletion is estimated in two ways. The &linear trend' method "rst assumes linear trends due to the anthropogenic in#uence of CFC emissions on total ozone, photodissociation rates and tropospheric species concentrations. Natural variations in stratospheric ozone such as the QBO and the solar cycle are excluded. The mean ozone seasonal cycle is used and the natural changes of ozone from year to year are included
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in the error estimation for the &linear trend' method. The minimum sensitivity of tropospheric composition to anthropogenic stratospheric ozone depletion is estimated by reducing the total ozone trends by their 2p errors. The use of the TOMS total ozone data to calculate photodissociation rates in the chemistry-transport model allows us to follow temporal changes in the tropospheric composition, and the assumption of linear trends in species concentrations is made at the end to determine trends. The e!ect of the photodissociation on tropospheric composition is studied by means of the 3D (103 latitude by 103 longitude) MOGUNTIA chemistry-transport model (Crutzen and Zimmermann, 1991; Dentener and Crutzen, 1993). The MOGUNTIA model is used in conjunction with monthly averaged transport (Krol, 1994), the advection-di!usion scheme is taken from Vested et al. (1992), and only background CH , CO, NO , and HO V V chemistry is considered. The calculated chemical composition of the atmosphere shows roughly the correct behaviour (Krol and Weele, 1997). The MOGUNTIA model uses the parameterization of photodissociation rates to represent the e!ects of stratospheric ozone absorption, clouds and surface re#ection (Krol and Weele, 1997). The TOMS total ozone data, the ISCCPC2 monthly averaged data on clouds, on snow/ice cover (Rossow and Schi!er, 1991) and the surface albedo data set (based on land-use data (Olson and Watts, 1982) and UV-A, UV-B average albedos (Feister and Grewe, 1995)) are used to calculate the photodissociation rates by means of the parameterization scheme. To study the sensitivity of tropospheric composition to changes in photodissociation rates, the MOGUNTIA model was run for a period of 15 years from 1978 to 1993 for three di!erent cases, the emissions being kept constant. First, the &no trend' MOGUNTIA model run was performed. The photodissociation rates were calculated via parameterization using the TOMS total ozone and ISCCP cloud data for one year from the beginning of the TOMS data record. Second, the &linear trend' MOGUNTIA model run was made. To the previously described photodissociation rates, calculated using 1 year of TOMS data, we added the linear trends in photodissociation rates of di!erent species due to stratospheric ozone depletion (see Chapter 3). These trends were calculated with the radiative transfer model every 2 h (MOGUNTIA time step) for the 15th of each month for latitude bands of 103. Third, the &reduced linear trend' MOGUNTIA model run was made. The aim was to determine the lower limits of tropospheric composition sensitivity to linear trends in total ozone due to the anthropogenic in#uence of CFC emissions. The stratospheric ozone trends are reduced by their 2p errors (De Winter-Sorkina, 2001) and the reduced trends in photodissociation rates are calculated. The di!erence between this run and the previous MOGUNTIA model run lies in the use of these reduced photodissociation trends. Fourth, the
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Fig. 4. Trends in OH, O , CO and NO in the boundary layer over the time period 1979}1992 due to trends in photodissociation rates. V Shaded areas represent the signi"cance area at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
&TOMS ozone' MOGUNTIA model run was performed. The whole 14 year record of TOMS total ozone data was used to calculate photodissociation rates via parameterization. The changes in OH, O , CO, NO and V CH in time were calculated relative to the &no trend' run and modelled by the seasonal cycle, the linear trend and the noise term in a linear regression "t. The trends in OH, O , CO and NO in the boundary V layer over the time period 1979}1992 due to the trends in photodissociation rates are shown in Fig. 4. Here &linear trend' and &no trend' MOGUNTIA model runs were used. Minimum trends in OH, O , CO and NO occur V red in the boundary layer over the time period 1979}1992 due to the trends in photodissociation rates reduced by their 2p errors are shown in Fig. 5. Here &reduced linear trend' and &no trend' MOGUNTIA model runs described above were used. Trends in OH, O , CO and NO in the boundary 6 layer over the time period 1979}1992 due to the trends in photodissociation rates calculated from the TOMS ozone dataset via parameterization are shown in Fig. 6. Here &TOMS ozone' and &no trend' MOGUNTIA model
runs were used. The signi"cance area at the 2p error level is shaded and was calculated by linear regression "t. The errors in trends describe the ability to represent the temporal dependence of species concentrations by seasonal term and the linear trend. In the case of trends determined by the &linear trend' method these errors are negligible due to the use of the mean ozone seasonal cycle. Note that errors in trends from the &TOMS ozone' model run do not include the long-term drift in TOMS data. Positive trends in tropospheric OH concentration (due to stratospheric ozone depletion) in the boundary layer of northern mid-latitudes of about #6}#8% per decade (#5}#7% per decade from the &TOMS ozone' run) in February-April were calculated. The OH trend at 453N reaches its maximum of #9.3$2.6% per decade in February (#7.3$3.1% per decade in March from the &TOMS ozone' run). Note that the in#uence of this trend on the global oxidation capacity is small due to low OH. The lower limits of OH trends (when the 2p errors are subtracted from the stratospheric ozone trends, see Fig. 5) of about #4}#6 % per decade were calculated in
R. De Winter-Sorkina / Atmospheric Environment 35 (2001) 1615}1625
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Fig. 5. Minimum trends in OH, O , CO and NO in the boundary layer over the time period 1979}1992 due to trends in V photodissociation rates reduced by their 2p errors.
February}April in the boundary layer of northern midlatitudes. The OH trends estimated from the &TOMS ozone' MOGUNTIA run are somewhat lower, but within the error limits of the OH trends determined from the &linear trend' run. The annual OH trend in the boundary layer at 453N is #4.3$2.8% per decade from the &linear trend' run and #4.0$1.0% per decade from the &TOMS ozone' run. A global trend in tropospheric OH concentration due to stratospheric ozone depletion amounting to #2.4$1.3% per decade (#2.9$0.6% per decade from the &TOMS ozone' run) was calculated. This can explain part of the global linear OH trend estimated from the methylchloroform measurements to be about #4.6$0.6% per decade (Krol et al., 1998). The temporal dependence of the change in global OH calculated from the &TOMS ozone' run agrees well with the ones presented by Bekki et al. (1994) and WMO (1999). The e!ect on tropospheric ozone concentration appears to be moderate. The negative trends in tropospheric ozone concentration (due to stratospheric ozone depletion) in the boundary layer at 553N reach their maximum of !2.2$1.2% per decade (!1.8$1.0% per
decade from the &TOMS ozone' run) in June. The lower limits of tropospheric ozone trends (when the 2p errors are subtracted from the stratospheric ozone trends) of about !0.5}!1% per decade were calculated in April}July for the boundary layer at 553N. The tropospheric ozone trends estimated from the &TOMS ozone' MOGUNTIA run are within the error limits of the tropospheric O trends determined from the &linear trend' run. The annual tropospheric ozone trend in the boundary layer at 553N is !1.0$0.7% per decade from the &linear trend' run and !0.8$0.3% per decade from the &TOMS ozone' run. The global trend in tropospheric O concentration due to stratospheric ozone depletion of !0.8$0.7% per decade (!0.7$0.1% per decade from the &TOMS ozone' run) was calculated at the 2p uncertainty level. Surface measurements of CO and CH from global networks have revealed a decline in growth rates during the last decade (WMO, 1999). The rate of OH increase, which is necessary to explain the change in the CO trend at northern mid-latitudes between 1980 and 1995, is estimated to be 6$3% per decade (Yurganov et al.,
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Fig. 6. Trends in OH, O , CO and NO in the boundary layer over the time period 1979}1992 due to trends in photodissociation rates V calculated from the TOMS ozone dataset via parameterization. Shaded areas represent the signi"cance area at the 2p error level. Black patches indicate missing TOMS data due to the polar night.
1999). The negative trends in tropospheric CO concentration (due to stratospheric ozone depletion) in the boundary layer at 453N reach their maximum of !2.2$1.1% per decade in May (!1.9$1.1% per decade in June from the &TOMS ozone' run). The lower limits of CO trends (when the 2p errors are subtracted from the stratospheric ozone trends) of about !0.5 to !1% per decade were calculated in February} June for the boundary layer at 453N. The CO trends estimated from the &TOMS ozone' MOGUNTIA run are within the error limits of the CO trends determined from the &linear trend' run. The annual CO trend in the boundary layer at 453N is !1.3$0.9% per decade from the &linear trend' run and !1.2$0.2% per decade from the &TOMS ozone' run. The global trend in tropospheric CO concentration due to stratospheric ozone depletion of !1.8$1.3% per decade (!2.0$0.2% per decade from the &TOMS ozone' run) was calculated. We might expect variation in the spatial NO chemV istry (Fuglestvedt et al., 1994) since increased OH levels will increase the transfer of NO to HNO , which is the main loss mechanism for NO . The negative trends in V
tropospheric NO concentration (due to stratospheric V ozone depletion) in the boundary layer at 553N reach their maximum of !1.7$0.4% per decade in April (!1.4$0.8% per decade in May from the &TOMS ozone' run). The lower limits of NO trends (when the V 2p errors are subtracted from the stratospheric ozone trends) of about !1% per decade were calculated in February}May for the boundary layer at 453N. The NO V trends estimated from the &TOMS ozone' MOGUNTIA run are within the error limits of the NO trends deterV mined from the &linear trend' run. The annual NO trend V in the boundary layer at 453N is !0.8$0.5% per decade from the &linear trend' run and !0.8$0.3% per decade from the &TOMS ozone' run. Urban measurements of NO in the United States show a downward trend of about !0.5% per year in the 1980s (Logan, 1994). Trend analyses of surface NO measurements in North-West V Europe over the period 1980}1994 (De Paus and Roemer, 1997) shows that concentrations of NO were V decreasing with 0.1}4.3% per year, except for UK, mainly as a result of decreasing emissions. No signi"cant global trend in tropospheric NO concentration due to V
R. De Winter-Sorkina / Atmospheric Environment 35 (2001) 1615}1625
stratospheric ozone depletion was found at a 95% con"dence level in our study, as a result of NO trends V decreasing with altitude. Due to increases in OH, the global level of CH is reduced. Because of the long lifetime of CH , the maximum e!ect on methane is delayed by a few years (Fuglestvedt et al., 1994). The negative trends in tropospheric CH concentration (due to stratospheric ozone depletion) in the boundary layer reach their maxima of !0.8$0.7% per decade at 753N in July and !0.9$0.8% per decade at 653S in December (!1.2$0.2% per decade at 753N in August and !1.3$0.2% per decade at 753S in December from the &TOMS ozone' run). The CH trends estimated from the &TOMS ozone' MO GUNTIA run are within the error limits of the CH trends determined from the &linear trend' run. The lower limits of CH trends (when the 2p errors are subtracted from the ozone trends) of about !0.08 to 0.1% per decade were calculated for the boundary layer at 453N. The annual CH trend in the boundary layer at 453N is !0.8$0.7% per decade from the &linear trend' run and !1.15$0.05% per decade from the &TOMS ozone' run. The moderate response of methane is explained by the lack of correlation between the changes in OH (largest at higher latitudes in spring) and the rate of methane oxidation, which increases with temperature (largest at lower latitudes) (WMO, 1999). The global CH trend (at 2p error level) of !0.8$0.7% per decade (!1.2$0.1% per decade from the &TOMS ozone' run) was found over the 1979}1992 time period.
5. Discussion and conclusions An ozone depletion climatology based on the TOMS/ Nimbus-7 (11/1978}04/1993) version 7 total ozone data and SAGE ozone trends together with ozonesonde data was used to estimate monthly trends in the daytime average photodissociation rate coe$cients of tropospheric species and their daily cycle due to stratospheric ozone depletion over 103 latitude bands. The minimum response of the tropospheric composition to anthropogenic stratospheric ozone depletion over this time period is estimated. The trends in tropospheric OH, O , CO, NO and CH distributions are estimated by the &linear V trend' and the &TOMS ozone' methods using the 3D MOGUNTIA model. The trends estimated from the &TOMS ozone' method are within the error limits of the trends determined from the &linear trend' method. This means that the &linear trend' approach can be used for estimating future scenarios when the yearly ozone variations are unknown. The error analyses of stratospheric ozone depletion and its impact on tropospheric trends shows that signi"cant changes have occurred in tropospheric densities at mid- and high latitudes of both hemispheres over the
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period 1979}1992. Global trends in tropospheric concentrations due to stratospheric ozone depletion are calculated by MOGUNTIA model of #2.4$1.3% per decade for OH, !1.8$1.3% per decade for CO and !0.8$0.7% per decade for O and CH when emis sions are kept constant. Thus, a small decrease in global tropospheric ozone was estimated due to stratospheric ozone reduction. However, the e!ect of stratospheric ozone depletion on tropospheric trends is small or comparable to other factors, especially the anthropogenic emission changes. Response of global OH concentrations to stratospheric ozone loss was found to be equal to a possible 10% NO emission increase and 85% of this V response was found to be equal to an estimated 6.5% CO decrease or a possible 10% tropical H O increase (Krol et al., 1998). The estimated 11% CH increase leads to a decrease of global OH of 55% of its positive response to stratospheric ozone depletion (Krol et al., 1998). However, it should be emphasized that the changes in the global atmospheric composition since 1978 are to a large extent unknown. Granier et al. (1996) calculated CO concentration responses to combined stratospheric ozone changes over period 1989/1990}1993, tropospheric temperature changes, a 10% decrease of industrial source of CO and a 25% reduction in biomass-burning (CO, NMHC, CH and NO emissions). Under this V scenario, the changes calculated for the 1990}1993 period are consistent with the changes observed at most CO measuring stations. However, total ozone and temperature change could be responsible for no more than 30% of the total observed change in CO (Granier et al., 1996). Trends in tropospheric ozone in the context of trends in emissions of NO were discussed by Logan (1994). V Reported trends in tropospheric ozone are highly variable, sparse and sometimes ambiguous. Thus, the impact of stratospheric ozone depletion on tropospheric ozone trends can be shadowed by the impact of emission trends. However, in regions were emission changes are playing a minor role, e!ects due to stratospheric ozone reduction can be indicated. For example, three Canadian ozonesonde stations show a long-term decrease in ozone since 1980, which is signi"cant only in winter and autumn and may re#ect dynamical variability and lower amounts of stratospheric ozone (Logan, 1994). Tropospheric ozone changes due to stratospheric ozone reduction are very sensitive to local NO concentrations (Liu V and Trainer, 1988). In general, tropospheric ozone trends due to stratospheric ozone depletion are too small compared to the current uncertainty of the long-term ozonesonde and surface ozone observations to be measured.
Acknowledgements Renata de Winter-Sorkina was supported by the European Union RIFTOZ project and by the Space Research
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Organisation of the Netherlands (SRON). The author thanks Maarten Krol for performing MOGUNTIA model runs, for helpful discussions and for comments on the manuscript, and is grateful to Peter Builtjes for his interest and support.
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