Lead Author: Mark R. Schoeberl
Co-authors: Anne R. Douglass John Gille James A. Gleason William R. Grose Charles H. Jackman Lamont Poole Steve Massie M. Pat McCormick Jim Miller Paul A. Newman Lamont Poole Richard B. Rood Gary Rottman Richard S. Stolarski Joe W. Waters
9.1.1.1 Location of the ozone layer and climatology
9.1.1.2 Ozone and UV - biological threat
9.1.1.3 Ozone and climate change
9.1.2.1 Polar ozone changes
9.1.2.2 Mid-latitude ozone loss
9.1.3.1 Chemical processes
9.1.3.2 Transport
9.1.3.3 Aerosols and Polar Stratospheric Clouds
9.1.3.3.1 Aerosols
9.1.3.3.2 Polar Stratospheric Clouds
9.1.3.4 Solar Ultraviolet and Energetic Particles
9.1.4.12-D Models
9.1.4.2 3-D Models
9.2.1.1 Interannual and long term variability of the stratospheric circulation
9.2.1.2 External influences (Solar and Particle effects)
9.2.1.3 Natural Aerosols and PSC's
9.2.2.1 Trends Chlorine Source Gases
9.2.2.1.1 Historical Trends in Chlorine Source Gases
9.2.2.1.2 Stratospheric Chlorine
9.2.2.1.3 Depletion of Ozone by Stratospheric Chlorine
9.2.2.2 Effects of aircraft exhaust
9.3.2.1Science questions
9.3.2.2Key trace gas measurements
9.4.1.1 Global limb temperature measurements
9.4.1.2 Higher horizontal resolution temperature profiles
Ozone is one of the most important trace species in the atmosphere. Ozone plays two critical roles, it removes most of the biologically harmful ultraviolet light before the light reaches the surface and it plays a essential role in setting up the temperature structure and therefore the radiative heating/cooling balance in the atmosphere, especially the stratosphere (the region between about 10 and 60 km)..
Ozone is mainly found in two regions of the atmosphere. Most of the ozone can be found in a layer between 10 and 60 km above the Earth's surface (Fig9.1.1-1). This ozone located in the stratosphere is known as the 'ozone layer.' Some ozone can also be found in the lower atmosphere (below 10 km), in the region known as the troposphere. Although chemically identical to stratospheric ozone, tropospheric ozone is quite distinct and geophysically different from stratospheric ozone and the science issues concerning tropospheric ozone will be discussed in Chapter 5.
Ozone is produced by the photolysis of molecular oxygen, O2. The oxygen atom, O, produced by this photolysis recombines with O2 to form ozone, O3. Ozone formation primarily occurs in the tropical upper stratosphere, where it is transported poleward and downward by the large scale Brewer-Dobson circulation. The global distribution of total ozone is shown in Fig 9.1.1-2. This figure represents the 13 year average of the total ozone measurements taken by the Nimbus-7 Total Ozone Mapping Spectrometer (TOMS) instrument.
The formation of ozone by the photolysis of molecular oxygen removes most of the light with wavelengths shorter than 200 nm. The wavelengths between 200 and 310 nm are removed by the photolysis of ozone itself. This photolysis of ozone in the stratosphere is the process by which most of the biologically damaging ultraviolet sunlight (UV-B) is filtered out.
As this filtering process occurs, the stratosphere is heated. This heating is responsible for the temperature structure of the stratosphere, where the temperature increases as the altitude increases. Without this filtering larger amounts of UV-B would reach the surface. Numerous studies have shown that excessive exposure to UV-B is harmful to plants, animals, and humans (WMO, 1992).
If ozone in the stratosphere were to be removed, the stratosphere would cool. How a cooler stratosphere affects radiative balance in the rest of the atmosphere as been the subject of many detailed studies. These studies have been re-analyzed and integrated in to the latest IPCC report, 'Climate Change 1994: Radiative Forcing of Climate Change.' The conclusions of that report state that stratospheric ozone loss leads to a 'small but non negligible offset to the total greenhouse forcing from CO2, N2O, CH4, CFCs...' It is ironic that the size of the negative radiative forcing from ozone loss is nearly equal to the positive radiative forcing from chlorofluorocarbons (CFCs), the source of the stratospheric ozone loss. The size of the radiative forcing due to stratospheric ozone loss has also been shown to be very sensitive to the profile shape assumed for that loss.
While the global amount of ozone is fairly constant, there are significant local, seasonal and long term changes. The cause of these changes are discussed in detail in the sections 9.1.3.
The seasonal ozone changes are basically determined by the winter-summer changes in the stratospheric circulation. Since ozone has a lifetime of weeks to months in the lower stratosphere, the amount of ozone can strongly vary due to transport by stratospheric wind systems. Since weather conditions in the stratosphere, like the troposphere, vary from year to year, there is also interannual variability in ozone amounts. Interseasonal changes in ozone are also linked to the 11 year solar cycle in UV output, and the amount of volcanic aerosols in the stratosphere. Changes in ozone have also been linked to anthropogenic pollutants especially the release of man made chemicals containing chlorine. In the section below we describe the more significant recent global changes in ozone observed by a variety of instruments.
The first ozone measurements in the Antarctic were made during the 1950s. A Dobson instrument was installed at Halley Bay in late 1956 in preparation for the International Geophysical Year in 1957. One of the first discoveries made by this instrument was that the seasonal cycle of ozone in the south polar region is very different from that which had been observed in the north. This was noted in a review article by Dobson (1966) which pointed out that its cause was a difference in the circulation patterns of the Antarctic relative to the Arctic. In the Arctic, the total ozone amount grew rapidly in the late winter and early spring to about 500 Dobson Units (DU). (A Dobson Unit is one milliatmosphere/cm^2 of pure ozone at STP.) In contrast, the Antarctic early springtime amounts remained near 300 DU.
The Dobson instrument at Halley Bay continued to make measurements each year. Farman et al. (1985) showed that the springtime ozone amounts over Halley Bay had declined from nearly 300 DU in the early 1960s to about 180 DU in the early to mid 1980s. This result has been confirmed at a number of other stations and shown using satellite data to occur over an area larger than the Antarctic continent (Stolarski, et al., 1986). Figure 9.1.2-1 shows an update of the Halley Bay data for October with TOMS and Backscatter Ultraviolet (BUV) satellite measurements included. These large ozone changes implied that losses must be taking place in the lower stratosphere wheremost of the ozone exists. This was shown to be true in a series of ozonesonde measurements (see e.g. Hofmann, et al., 1989 and Figure 9.1.2-2. More recent sonde measurements have shown instances of near-zero concentrations of ozone over a 5-km altitude range (Hofmann, et al., 1994). Aircraft measurements (Proffitt, et al., 1989) and satellite measurements (McCormick et al., 1988) confirm and show further details of these ozone changes.
Ozone changes in the Arctic have not been nearly so dramatic (until 1996). Because of the larger seasonal and interannual meteorological variations in the northern hemisphere, trends in total ozone are more difficult to detect. Both the austral and boreal winter stratospheres are dominated by a strong circumpolar jet called the polar vortex. Large wave disturbances in the northern hemisphere distort the vortex and cause its earlier breakup with an uncertain timing. Analyses show that the trend is less than 10% per decade and is spread through much of the midlatitudes during winter and spring (see next section for more detail). Recent measurements from UARS (Manney et al., 1995) show that polar ozone in the northern hemisphere varies from year to year dependent on the coldness and persistence of the vortex in a manner consistent with chemical depletion of ozone by chlorine.
In 1996, an extremely cold and persistent Arctic vortex formed. UARS measured high levels of chlorine from December up to early March. Estimates of ozone loss, as of this writing, approach 50%. Although the Arctic stratosphere has been unusually cold, the temperature conditions are not outside of the range of climatology.
Midlatitude ozone loss estimates must be extracted from long time series using statistical models (see e.g. Ozone Trends Panel, 1990). The longest time series of total ozone data is from Arosa in Switzerland. This time series, which dates back to 1926, is shown in Figure 9.1.2-3. The Arosa data show a relatively constant amount of ozone for over 4 decades and a decrease in the last decade and a half. Analysis of a more extensive network of 30+ stations which have been in operation for about 35 years show negative trends over the last 1.5 decades, especially in the winter and early spring (see e.g. Reinsel, et al., 1994).
High quality global satellite data records begin on November1978 with the launch of NIMBUS 7 SBUV and TOMS instruments. These data show midlatitude trends in the northern hemisphere which are largest in the winter and early spring, peaking at about 6-8% per decade at 40-50N in February (see e.g. Randel and Cobb, 1994; Hollandsworth, et al., 1995). These satellite trends are in the process of being updated with a version 7 algorithm for the TOMS and SBUV instruments.
Changes in the profile of ozone with altitude can be deduced from sonde data or from the SAGE satellite measurements. Analyses of sonde data (e.g. Logan, 1994) show ozone decreases between the tropopause and about 24 km altitude. Analyses of SAGE data (McCormick and Larson, 1992) show larger decreases than those derived from sondes. SAGE results show negative ozone trends in the lower stratosphere in the tropics. Column ozone changes deduced from SBUV and TOMS show only small downward trends. Hollandsworth, et al., (1995) used SBUV profile and total ozone trends to deduce that ozone in the tropics below 32 mbar has increased slightly over the last decade. The resolution of the uncertainty in the magnitude of lower stratospheric and upper tropospheric ozone trends is an important measurement and analysis issue for the coming years.
Ozone is being continuously created and destroyed by the action of ultraviolet sunlight. The overall amount of ozone in the global stratosphere is determined by the magnitude of the production and loss processes and by the rate at which air is transported from regions of net production to those of net loss.
Production of ozone requires the breaking of an O2 bond with the extra or "odd" oxygen atom attaching to another O2 to form O3. This most frequently occurs via the photodissociation of O2 by solar ultraviolet radiation. In the lower stratosphere and troposphere ozone can also be produced by photochemical smog-like reactions. In these reactions H or CH3 or higher hydrocarbon radicals attach to an O2 forming HO2 or CH3O2, etc. which then react with NO. This reaction breaks the O2 bond by forming NO2 (which is really ONO). When NO2 is photolyzed an O atom is formed which reacts with O2 to form O3.
Loss of ozone occurs when an O atom reacts with O3 to reform the O2 bond. More importantly, this loss process is catalyzed by the oxides of hydrogen, nitrogen, chlorine, and bromine. These oxides are produced in the stratosphere from long-lived, unreactive molecules released at the surface of the earth. The major source molecules for HOx are methane (CH4) and water vapor (H2O). The main source of NOx is nitrous oxide (N2O). The major sources of chlorine are industrially produced chlorofluorocarbons (such as CFC-11 which is CFCl3 and CFC-12 which is CF2Cl2) and naturally occurring methyl chloride (CH3Cl). The major sources of bromine are methyl bromide (CH3Br) and the halons (CF3Br and CF2ClBr). These source molecules are transported to the stratosphere where they react or are photodissociated to produce the catalytically active oxide radicals.
The catalytic efficiency of hydrogen, nitrogen, chlorine, and bromine oxides is determined by a set of interlocking reactions which convert the active oxides to catalytically inactive temporary reservoirs, such as HNO3, HCl, ClONO2, H2O, HOCl, HOBr, and BrONO2, and vice versa. In the lower stratosphere, the balance between catalytic oxides and temporary reservoirs is strongly affected by reactions on the surfaces of stratospheric aerosols. The balance is even more profoundly affected in the polar winter by reactions on the surface of polar stratospheric cloud particles. In the early spring, the chlorine balance is shifted to almost 100% ClOx (Brune et al., 1989; Waters et al., 1994). This shift in the chemical balance results in a large calculated chemical sensitivity of ozone towards chlorine perturbations and a relatively small calculated sensitivity of ozone towards nitrogen oxide perturbations.
Although the basic outline of the chemistry controlling stratospheric ozone is now known, many important aspects of the problem remain to be solved. The primary difference between the northern and southern hemispheric polar ozone loss regions appears to be a result of the "denitrification" that occurs in the Antarctic winter. Denitrification means the removal of nitrogen oxides and HNO3 by large particles which fall into the troposphere. Denitrification takes place when temperatures are cold enough to form large stratospheric ice crystals.
When springtime comes there are no nitrogen oxides to convert ClOx to ClONO2 and slow down the rate of ozone depletion. There is some evidence for denitrification when temperatures are not cold enough to form ice crystals. Under those conditions the mechanism for denitrification is not completely understood.
Much of the currently observed ozone interannual variability in the stratosphere is controlled by dynamical processes. In particular, this variability is driven by such processes as the quasi-biennial oscillation (QBO), El Nino- Southern Oscillation, tropospheric weather systems which extend into the stratosphere, and long term fluctuations in planetary wave activity. The annual cycle of total ozone is largely driven by transport effects. As discussed in section 9.1.1, relatively low values of ozone are observed in the tropics and high values are observed in the extratropics. These low tropical ozone values occur in spite of the large ozone production rates in the tropics. If ozone production were precisely balanced by ozone loss everywhere, total ozone would have extremely high values in the tropics. The observed tropical low values result from vertical advection of low ozone air from the tropical troposphere into the tropical stratosphere, and the subsequent transport of this air poleward and downward into the extratropics and polar regions. This advective circulation is known as the Brewer-Dobson circulation.
The redistribution of ozone from the production region at low latitudes to extratropical latitudes is modulated by a variety of processes. Foremost among these processes is the annual cycle in the circulation. It is now recognized that the Brewer-Dobson circulation is primarily controlled by large scale waves in the winter stratosphere. As these waves propagate through the westerly winds that dominate the winter stratosphere, they drive the temperatures away from radiative equilibrium, which leads to a poleward and downward transport circulation. The large scale waves in the winter upper-stratosphere also produces lifting in the tropics. Since the lifetime of ozone increases with pressure, the poleward downward circulation causes ozone to accumulate in the lower stratosphere over the course of the winter. Since the large scale waves are not present in the summer, the poleward and downward circulation is significantly weakened and ozone amounts which have built up during winter begin to decrease due photochemistry.
It is now known that knowledge of stratospheric aerosols and polar stratospheric clouds (PSCs) are very important to our understanding of stratospheric ozone. The surface of aerosols and PSCs are sites for heterogeneous reactions which can convert chlorine from reservoir to radical forms. Likewise, radical nitrogen forms can be sequestered as nitric acid to shift the chemical loss process (WMO, 1995).
The long-term stratospheric aerosol record reveals at least three components: episodic volcanic enhancements, PSCs and clouds just above the tropical tropopause, and a background aerosol level. At normal stratospheric temperatures, aerosols are most likely super cooled solution droplets of H2SO4-H2O, with an acid weight fraction of 55 to 80%. The primary source of stratospheric aerosols is volcanic eruptions that are strong enough to inject SO2 buoyantly into the stratosphere. Aerosol sizes range from hundredths of a micrometer to several micrometers. Although there is some variability, especially just after a volcanic eruption, a log-normal size distribution of spherical particles appears to aptly describe the aerosol. Just after an eruption, the size distribution becomes bimodal, and some particles are nonspherical because of the addition of crustal material. After an eruption, the SO2 is converted to H2SO4 with a time scale of about 30 days. Subsequently stratospheric sulfuric acid aerosols, aerosol loading decreases with an e-folding time of 9 to 12 months, although this appears quite variable with altitude and latitude. The decay is most likely due to a combination of sedimentation, subsidence and exchange through tropopause folds.
The net effect of this post-volcanic dispersion and natural cleansing is a greatly enhanced aerosol concentration in the upper troposphere after a major eruption, especially poleward of about 30°latitude. Except immediately after an eruption, stratospheric aerosol droplets tend to be concentrated into 3 distinct latitudinal bands, one over the equatorial region (to 30°) and the other over each high latitude region, 50 to 90°N and S. Following a low latitude eruption, aerosol is dispersed into both hemispheres, whereas following a mid-to-high latitude eruption, aerosols tend to stay primarily in the hemisphere of the eruption. Potential sources of a background aerosol component include OCS from the oceans, low level SO2 emissions from volcanoes, and various anthropogenic sources, including industrial and aircraft emissions. Also, it is not clear whether there is an upward trend in this background aerosol, as has been hypothesized and linked to increasing aircraft emissions, since any increase may be due to incomplete removal of past volcanic aerosol.
Stratospheric aerosol loading in 1979 was approximately 0.5 x 1012g (0.5MT), thought to be representative of background aerosol conditions. The present status of the aerosol is one of enhancement due to the June 1991 eruption of Pinatubo (15.1 degrees N, 120.4 degrees E), which produced on the order of 30 x 1012g (30MT) of new aerosol in the stratosphere, about 3 times that of the 1982 eruption of El Chichon. This perturbation appears to be the largest of the century, perhaps largest since the 1883 eruption of Krakatoa. By early 1993, stratospheric loading decreased to approximately 13 MT, about equal to the peak loading values after El Chichon [McCormick et al. ,1995]. Measurements in 1995 show that the aerosol levels are approaching background levels.
The interannual variability in PSC sightings has been addressed by Poole and Pitts (1994), who analyzed more than a decade of data from the spaceborne Stratospheric Aerosol Measurement (SAM) II sensor (Fig. 9.1.3.3.2-1). They found noticeable variability in PSC sightings in the Antarctic from year to year, even though the southern polar vortex is typically quite stable and long-lived. This variability was found to occur late in the season and can be explained qualitatively by temperature differences.
Poole and Pitts found even more year-to-year variability in SAM II Arctic PSC sighting probabilities. This was expected since the characteristics and longevity of the northern polar vortex vary greatly from one year to the next. The year-to-year variability in Arctic sighting probabilities can also be explained qualitatively by differences in temperature; e.g., zonal mean lower stratospheric temperatures in February 1988 were as much as 20 K colder than those one year earlier.
Since ozone formation is fundamentally linked to the levels of ultraviolet radiation reaching the earth, natural variations in that radiation must be understood in order to detect trends. The ultraviolet comprises only one to two percent of the total solar radiation, but it displays considerably more variation than the longer wavelength visible radiation. For example, from 1986 to 1990 the solar UV increased with onset of the 11-year solar cycle and resulted in an increase of global total ozone of almost 2%. This natural increase in ozone is comparable to the suspected anthropogenic decrease and needs to be understood in order to totally separate the anthropogenic decrease from this natural change. Studies of total ozone trends typically subtract solar cycle and other natural changes from the total ozone record in trend resolution (see WMO 1991; WMO 1995; Stolarski et al., 1991; and Hood and McCormick, 1992), thus more quantitative knowledge of this natural solar cycle-induced total ozone change would be especially valuable.
Changes in energetic particle flux from the Sun penetrate into the middle atmosphere and may also drive the natural ozone variations. A series of solar flares in 1989 spewed solar particles into the earth's polar cap regions (greater than 60° geomagnetic latitude) and led to polar ozone depletion (Jackman et al., 1993). Further work on the very large solar particle events (SPEs) of October 1989 have predicted ozone depletions lasting for several months after the SPEs (Reid et al., 1991; Jackman et al. 1995). Although SPEs of this magnitude occur infrequently (only two have been observed in the past 25 years), they need to be understood more completely to be able to separate natural from anthropogenic ozone effects.
Relativistic electron precipitations (REPs) have been predicted to contribute substantially to the odd nitrogen budget of the stratosphere and, therefore, predicted to play a large role in controlling ozone in this region (Callis et al., 1991a, 1991b). Another investigation (Aikin, 1992) has failed to find any REP-caused ozone depletion. Further work (Gaines, et al. 1995) determined that REPs in May 1992, the largest measured relativistic electron flux precipitating in the atmosphere between October 1991 and July 1994, added only about 0.5 to 1% of the global annual source of odd nitrogen to the stratosphere and mesosphere. The actual importance of REPs in regulating ozone is thus not well understood nor characterized, and further work on REPs is required to thoroughly determine their importance regarding modulation of stratospheric ozone.
Models of the stratosphere provide the only means to attempt quantitative prediction of global change, or to evaluate the impact of natural or anthropogenic changes in composition on the stratospheric ozone and climate. In addition, models provide a means to integrate observations and theory, to provide tests of mechanisms for chemical, dynamical or radiative changes, and to enable interpretation of observations from different platforms.
Two-dimensional (2D) models are used by several research groups. The models predict the behavior of ozone and other trace gases in reasonable agreement with measurements (WMO 1991; WMO 1995). Because of these favorable comparisons to measurements, these models have been utilized recently in many atmospheric studies, for example: (1) the response of the middle atmosphere due to solar variability was studied by Brasseur (1993), Huang and Brasseur (1993), and Fleming et al. (1995); (2) the influence of Mt Pinatubo eruption on the stratosphere was studied by Kinnison et al. (1994) and Tie et al. (1994); and (3) the effects of the proposed stratospheric aircraft on atmospheric constituents were studied by Pitari et al. (1993), Weisenstein et al. (1993), Considine et al. (1994), and Considine et al. (1995).
These 2D models have also been used to produce multi-year simulations of the response of stratospheric ozone to perturbations of the source gases such as chlorofluorocarbons from which chlorine radicals are produced (WMO 1991; WMO, 1995). An outstanding issue regarding simulations of the stratospheric ozone response to chlorine increases is the lack of ability of 2D models to accurately predict the ozone trend in the middle and high northern latitudes over the 1980 to 1990 time period. Since the 2D models predict a smaller trend than observed, it is believed that the models do not adequately model all of the relevant processes and thus require further development.
The three dimensional (or general circulation) model with full interaction between chemical, dynamical, and radiative processes remains elusive. The present generation of general circulation models generate unrealistic temperature fields which, in turn, alter the photochemistry. The unrealistic temperatures are related to problems with the model transport circulation. For example, the polar regions are persistently cold in general circulation models which suggests that there is insufficient adiabatic heating (or descent) in the winter polar region. Correspondingly, there will be insufficient ascent in the tropics which weakens the transport from the troposphere into the stratosphere.
Subtle changes in the general circulation of the atmosphere in 3-D models can alter and distort the chemical feedbacks. For example, Rasch et al. [1995] report on a two year simulation using version 2 of the NCAR Middle Atmosphere Community Climate Model (MACCM2). A chemical scheme for 24 reactive species or families is run as part of this simulation. This model is partially coupled in that the water vapor predicted by MACCM2 is connected to the chemical source of water through oxidation of methane. Prescribed ozone is used in the radiative calculation. In this simulation, the calculated upper stratospheric ozone is substantially lower than observations; much of the difference is attributed to the lower CH4, compared to observations (by the UARS Halogen Occultation Experiment (HALOE)). This bias leads to excessive ClO and excessive destruction of O3. In effect, the error in the this long lived trace gas which results from the weak transport circulation leads to noticeable errors in ozone.
The difficulties described above show why most 3-D modeling efforts have focused on "off-line" calculations, i.e., calculations of chemistry and transport (CTM's) in which the wind and temperature fields are input from a general circulation model [e.g., Eckman et al., 1995] or from a data assimilation system [e.g., Rood et al., 1991; Lefevre et al., 1994]. For either approach, there are computational advantages, as the same set of winds and temperatures are used many times. Furthermore, the effects of modifications to the chemical scheme can be isolated, and their effects understood without the complications caused by feedback processes. A further advantage of the use of assimilated winds and temperatures is that the results of constituent simulations may be compared directly with observations with no temperature biases such as those found in general circulation models. This is particularly important for the study of processes which have a temperature threshold, such as heterogeneous reactions on polar stratospheric cloud surfaces. The most information is gleaned when the model is sampled in a manner consistent with the satellite sampling [Geller et al., 1993].
The "off-line" approach has been used successfully for many years and are used to test chemical and transport mechanisms, as well as to interpret observations. These tests include: 1) assessment of the importance of transport of air with high levels of reactive chlorine to middle latitudes (Douglass et al., 1991); 2) assessment of the rate of ozone loss within the northern hemisphere vortex, and identification of the variables to which the calculation is sensitive (Chipperfield et al., 1993); 3) determination of the importance of upper tropospheric synoptic scale systems on the vortex temperature, as well as their influence on the transport and mixing of air which has experienced temperatures cold enough for polar stratospheric cloud formation (Douglass et al., 1993); 4) examination of the impact of ozone transport following the breakup of the Antarctic polar vortex on the global ozone budget (the "ozone dilution" effect).
These 3D studies provide a picture of the important physical processes which control polar ozone loss. However, because of computer resource restrictions, it is not yet possible to make full 3D model long range predictions, including possible influence of the ozone loss on lower stratospheric temperature and climate. For example, future temperature changes may have a significant impact on the northern hemisphere vortex. For example, Austin et al. [1992] discuss the probability forthe formation of an Arctic ozone hole in an idealized numerical simulation. The full three-dimensional model, with all relevant chemical, dynamical and radiative processes and feedbacks among them has yet to be developed.
Changes in the ozone layer can be divided into two categories: natural changes and man-made changes. Separating these components is the goal of much ozone and trace gas research. Since ozone can be transported by stratospheric winds, there is significant interannual variability in column ozone amounts. Ozone is likewise influenced by aerosol amounts (through heterogeneous chemistry) (Solomon et al., 1996), the formation of nitrogen radicals associated with high energy particles, and variations in the ultra-violet radiation from the sun. Man-made changes generally include increased chlorine and hydrogen amounts from industrial gases and increased aerosols and nitrogen radicals from airplane exhaust. Many of our current scientific issues and future measurement needs center around the interaction of the ozone layer with these pollutants and separating natural changes in the ozone layer from man-made processes.
Because the stratospheric circulation is strongly dependent on the dissipation of large scale waves in the stratosphere, interannual variability of the wave amplitudes has an important impact on ozone transport (see section 9.1.3.2). Winds and temperatures derived from 3-D general circulation models and assimilation models include such interannual variability and can be used to assess the impact on ozone transport. Two dimensional models can incorporate prescribed variability to simulate interannual ozone transport (see section 9.1.4.1). Accurate assessment of the of large scale waves and transport the circulation is necessary for understanding the variability of ozone trends.
One of the failures of the 3D models is an adequate simulation of the QBO. The QBO is a 24-30 month oscillation of the zonal wind in the tropical lower stratosphere that is driven by tropical waves. The QBO affects the stratospheric temperature distribution and produces a secondary circulation which transports trace gases and aerosols. For ozone, these variations can reach values of 5-10 DU in the tropics, and values of 10-20 DU in the extratropics.
The QBO provides one of the largest components of the interannual variability of the column ozone values. Because the geostrophic relationship breaks down in the tropics, direct tropical wind measurements are critical to precisely measuring the QBO and for understanding the affects of the QBO on the circulation. Data sparse regions, and infrequent sampling of wind fields all preclude good quantitative studies of the tropical circulation and its effect on ozone.
As discussed in section 9.1.3.4 solar ultraviolet radiation and precipitating energetic particles can strongly influence ozone amounts. In order to understand the anthropogenic changes in ozone, we must maintain reliable measurements of the solar ultraviolet input to the middle atmosphere. Solar variations in the UV produce ozone changes on the same order of magnitude as the current observed midlatitude changes. Proxies for the UV changes have been historically used to estimate the response of ozone to solar ultraviolet changes. With direct measurements from UARS, these proxies have been shown to inadequately represent changes in ultraviolet flux.
Particle events generate NOx compounds which catalytically destroy ozone, but these events tend to be confined to the upper stratosphere. Large events, which tend to be more episodic, may affect polar ozone at lower levels. The impact of NOx generation through particle precipitation on the natural ozone layer is a major scientific question.
As discussed in section 9.1.3.3, aerosols and PSCs are believed to play a major (although indirect) role in ozone loss. Irregular volcanic inputs of SO2 with the subsequent formation of sulfate aerosols have an impact on the ozone layer. There is some evidence suggesting that increasing amounts of background aerosols are a result of subsonic aircraft emissions in the lower stratosphere. A major scientific question is whether the background amounts of these aerosols are increasing, and if so, determining their origin. Monitoring the aerosol amounts within the stratosphere and determining their trend is a primary measurement requirement to understand ozone loss.
During the 1980's it became apparent that aerosols play an important role in the chemistry of the stratosphere. Observations of large decreases in ozone over Antarctica during the southern hemisphere spring were not accounted for by theory, until several researchers hypothesized that heterogeneous reactions on PSCs might be converting inactive chlorine compounds into reactive forms (Solomon et al., 1986; McElroy et al., 1986; Toon et al., 1986).In a similar fashion to PSCs, heterogeneous reactions upon sulfuric acid droplets at mid-latitudes converts N2O5 into HNO3 and shifts the ratio of HNO3 to NO2 normally present in the stratosphere. Throughout the stratosphere, reactions on and inside aerosol particles are therefore important.
To understand the effectiveness of the heterogeneous (gas phase/aerosol phase) reactions, it is important to know: a) the temperature of an aerosol particle b) the surface and volume densities of the aerosol particles, which are derived from the aerosol extinction, and a knowledge of the size distribution, c) the composition (the mixing ratios of H2O, H2SO4 and HNO3 in ppbv) and phase (liquid/solid/amorphous solid solution) of the aerosol particles, d) the concentration of the reactants in the aerosol (e.g. the concentration of HCl), and e) the duration of time over which the heterogeneous reactions occur. A theoretical framework, by which heterogeneous rates of reaction are quantified, is given in Hanson et al. (1994).
An important research goal is the ability to observe the yearly episodes of ozone loss in the polar regions (e.g. the Antarctic ozone hole), to measure this loss as reservoir chlorine levels change with time, and to be able to relate the changes in observed ozone to a quantitative understanding of heterogeneous processes.
In principle, one should be able to identify the composition of stratospheric aerosol from multi-wavelength extinction data. Multi-wavelength observations of mid- latitude sulfuric acid droplets have an extensive history. Observations of El Chichon aerosol (Pollack et al., l991), post El Chichon aerosol (Osborn et al., 1989; Oberbeck et al., 1989), and of Mt. Pinatubo aerosol (Grainger et al., 1993; Massie et al., 1994; and Rinsland et al., l994) yield spectral data consistent with theoretical expectation. Analysis of multi-wavelength observations of PSCs is a developing research topic. Recent attempts to use spectra to determine PSC composition is illustrated by Toon and Tolbert (1995).
Several years ago, ice and NAT (nitric acid trihydrate) particles were thought to be the primary composition of PSCs. Recent studies have shown that some PSC particles are liquid (the ternarysolution of HNO3/H2O/H2SO4), and not that of crystalline NAT (Carslaw et al., l994; Drdla et al., l994). As additional laboratory cold-temperature measurements of the indices of refraction of PSC composition candidates become available, the ability to classify PSC composition from spectra will improve.
Although PSCs are now known to be the instrumental in polar ozone loss, their amounts and types must be monitored. The major difference between the Antarctic ozone depletion and the less severe Arctic depletion appears to be the result of a lack of denitrification in the Arctic (Schoeberl et al., 1993). Fundamentally, denitrification is a function of temperature and the size of PSC's. Above frost point the PSC size is generally too small to precipitate nitric acid from the stratosphere. If temperatures reach frost point, larger PSC's form which are able to remove nitrogen acid from the lower stratosphere. The temperature history of the air parcel may play an important role in the PSC size distribution as well (e.g. Murphy and Gary, 1995). Photolysis of the nitric acid is key to halting the ozone depletion during winter.
With the increase of greenhouse gases, the stratosphere is expected to cool and thus increase the probability of PSC formation as well as increase the surface area and heterogeneous reaction rates on sulfate aerosols. Preliminary studies (Austin et al, 1992) suggest greenhouse gas increase could have a major role in polar ozone depletion through increased probability of PSC formation. Monitoring stratospheric aerosol loading and PSC amounts is critical for understanding ozone loss.
Man-made changes in ozone mostly arise from the manufacture of unreactive chlorine containing compounds such as the chlorofluorocarbons (chlorine source gases). These compounds reach stratospheric altitudes where photolysis by ultraviolet radiation releases chlorine with subsequent destruction of ozone through catalytic cycles. Aviation also has an impact on ozone through the release of nitrogen radicals in aircraft exhaust. Both of these anthropogenic effects are discussed below.
As mentioned earlier, chlorine source gases and their respective trends are the major drivers behind decreases in stratospheric ozone. A comprehensive discussion of chlorine source gases is contained in WMO (1995). This report may be consulted for more detail and appropriate references.
All chlorine in the stratosphere comes from tropospheric sources, predominantly the man-made chlorofluorocarbons (CFCs) and chlorocarbons. The man-made sources account for about 7/8th of the total stratospheric chlorine. CFCs are currently being phased out for the hydrochlorofluorocarbons (HCFCs) Extensive measurements of the chlorofluorocarbons CFC-11 (CCl3F), CFC-12 (CCl2F2), and CFC-113 (CCL2FCClF2) have indicated a steady increase in their tropospheric mixing ratios for more than a decade. Most recent data suggest that the growth rate for these species has begun to decrease. Measurements taken from Tasmania suggest that levels of the important chlorocarbon CCl4 in the troposphere are also decreasing.
As HCFC's are introduced as substitutes for CFC's, it may be expected that their mixing ratios in the troposphere will increase well into the next century. HCFC-22 (CHClF2) data show a near-linear growth rate in recent years. HCFC-141b and HCFC-142b have been available onlyrecently as CFC replacements. These species are clearly increasing in the troposphere, but further data is required to get reliable growth rates for long-term studies.
CH3CCl3 data also indicate a reduced growth rate that is a result of recently reduced emissions, but also possibly due in part to increasing hydroxyl (OH) levels. Data for dichloromethane (CH2Cl2), methyl chloride (CH3Cl), and chloroform (CHCl3) currently exhibit no long-term trends. Continued tropospheric measurements of these gases are required to estimate ozone depletion potential.
The most comprehensive suite of simultaneous measurements of chlorine constituents in the stratosphere include the ATMOS (Atmospheric Trace Molecule Spectroscopy) experiments of 1985, 1992, and 1993 and the AASE II (Airborne Arctic Stratospheric Expedition II) measurements of 1991, 1992. The data from these missions have provided invaluable information on the stratospheric chlorine burden and the partitioning among the various chlorine species.
Based upon the 1985 ATMOS data, Zander et al. (1992) determined a total stratospheric chlorine level of 2.55 + 0.28 ppbv. Further, they concluded that above 50 km most of the inorganic chlorine was in the form of hydrogen chloride (HCl) and that the partitioning of the chlorine among sources, sinks, and reservoir species was consistent with that level of total chlorine.
>From the 1992 ATMOS flights, total stratospheric chlorine (based upon HCl data above 50 km) was estimated to be 3.4 +/- 0.3 ppbv, an increase of approximately 35% in seven years (Gunson et al., 1994). This increase is consistent with that predicted by models (e. g. WMO, 1992). Schauffler et al. (1993) inferred total chlorine levels of 3.50 +/- 0.06 ppbv from the AASE II data near the tropopause, a value which is in excellent agreement with the 1992 ATMOS values.
Recent HCl data (55 km) from the Halogen Occultation Experiment (HALOE) on UARS (Russell et al., 1996) reveal a trend of 3.4 % per year in HCl for the period from UARS launch in 1991 to early 1995 (Figure 9.2.2.1.2-1).
Of the total stratospheric burden, only about 0.5 ppbv is estimated to arise from natural sources in the troposphere (WMO, 1995), but these estimates have yet to be confirmed by direct or remote observations. HCl emissions from major volcanic eruptions (El Chichon, 1982 and Mt. Pinatubo, 1991) provided negligible perturbations to the levels of HCl in the stratosphere (Mankin and Coffey, 1984; Wallace and Livingston, 1992; and Mankin et al., 1992).
Estimates of the severity of ozone depletion in the future can only be determined by atmospheric model simulations. The level of confidence in these models is based upon their ability to simulate present atmospheric distributions and their ability to simulate recent (decadal) trends. A discussion of the strengths and weaknesses of current assessment models is contained in WMO (1995) and section 9.1.4.1 above.
Model simulations of ozone change spanning the period 1980 to 2050 were conducted as part of the WMO (1995) assessment process. Two scenarios were adopted for the assessment studies: 1) the emissions of halocarbons follow the guidelines in the Amendments to the Montreal Protocol ,Scenario I; and 2) partial compliance with the guidelines, Scenario II (see WMO (1994) for specific details of the scenarios and models).
Figure 9.2.2.1.3-1 summarizes the results of the model calculations for Scenario I. This figure shows the percent change (relative to 1980) in the ozone column at 50N in March for each of the models participating in the assessment. Decreases of up to approximately 6.5 % are seen to occur just prior to 2000. The recovery to 1980 levels varies widely for the different models, from as early as 2020 to well past 2050. The individual models all showed reasonable agreement among themselves for the present day ozone distributions, but begin to differ substantially as the atmosphere is perturbed away from its existing state by increasing levels of nitrous oxide, methane, halocarbons, and other influences.
Uncertainties in the absolute levels of depletion predicted by the models are difficult to evaluate for these long-term scenario calculations. The trends in the source gases are changing and the trends in the stratospheric reservoir gases, which are dependent on transport into the stratosphere, will respond. Thus, measurements of the chlorine source and stratospheric reservoir gases must be made to test models against observations. Critical gases in the suite of required measurements are the reservoirs HCL and CLONO2. The predictive capability of these assessment models directly rests on additional measurements of chlorine source gases, reservoir gases and gases, which are sensitive to transport processes.
Long-lived source gases (e.g. N2O, CH4) are unreactive in the troposphere and hence can enter the stratosphere at the ambient tropospheric concentrations. In the stratosphere, these gases undergo photolysis or react with radicals to release their potential ozone destroying catalytic agents. In contrast, aircraft, flying in the stratosphere, will directly inject catalytic agents into the stratosphere. The primary agents for potential ozone change which have been considered in studies of aircraft exhaust are the nitrogen oxides (NOx) and water vapor (which leads to HOx). Now that heterogeneous reactions on background aerosols and PSCs are known to play an important role in the ozone balance of the stratosphere, the evaluation of the effects on ozone of NOx from supersonic aircraft flying in the stratosphere has changed significantly.
The impact on column ozone of a fleet of supersonic transports (now referred to as High Speed Civil Transports or HSCTs) is now calculated to be of the order of 1% or less. This is comparable to the calculated effects of HOx from the water vapor in the HSCT exhaust (NASA RP-1333, 1993). An important possibility is that the sulfur in the exhaust will lead to the generation of numerous small particles which will add to the aerosol surface area. An increase in surface area will enhance the conversion of chlorine from its reservoirs to ClOx and thus could lead to an increased loss rate for ozone. Another possibility is that the other condensibles in the exhaust, water vapor and nitric acid (from NOx), could impact the formation or duration of PSCs. Initial calculations show this effect to be small (Considine et al., 1995) and transport studies show that injection into the polar vortex unlikely (Sparling et al., 1995), but there is still uncertainty about what will happen as the stratosphere cools with increasing CO2 concentrations.
All of the chemical effects of HSCT exhaust depend on how much of the exhaust products accumulate in the stratosphere and where they accumulate. The same is true for the exhaust of the subsonic fleet which is released in the upper troposphere and lower stratosphere. The two major potential effects of the subsonic fleet of aircraft are ozone increase due to the smog-like photochemistry of NOx and contrail, CO2 increase due to fuel consumption, and cirrus cloud formation from the water vapor. The importance of aircraft NOx to ozone generation in the uppertroposphere and lower stratosphere is hotly debated. The role of aircraft as a source of upper tropospheric NOx is uncertain when compared to the NOx sources due to lightning, stratospheric intrusions, and the lofting of ground level pollution in cumulus clouds. Also uncertain is whether heterogeneous chemistry on ice crystals plays a significant role in the NOx budget.
Understanding the impact of supersonic and subsonic aircraft exhaust on the stratospheric chemical balance is a complex problem. Knowledge of meteorological conditions is required to compute exhaust dispersion. Knowledge of aerosol chemistry is required to understand the aerosol formation process (from sulfur in fuels) and its impact on the background conditions. Finally, a good understanding of the lower stratosphere chemistry is required to understand the direct impact of the NOx pollutants.
The investment by the scientific community in instrument and model development has produced a significant increase in our understanding of stratospheric chemical and dynamical processes. Although some fundamental questions of ozone loss have been answered, new questions have arisen. For example, the long term response of the ozone layer to natural fluctuations (QBO, El Nino, volcanoes) is still not well understood (section 9.2.1.1). The secular decrease in ozone following the eruption of Mt. Pinatubo was clearly associated with aerosol loading of the stratosphere - but the near one year delay in the appearance of maximum ozone loss is still not explained. More fundamentally, the midlatitude trend in column ozone loss reported by Stolarski et al. (1991) is still not explained (although it is probably connected with the increase in stratospheric chlorine and the stratospheric chemistry associated with aerosols, see section 9.1.3.3). Our understanding of the more subtle chemical processes is still quite incomplete which increases our uncertainty in the forecast predictions.
Under the Atmospheric Effects of Aviation Program (AEAP) the impact of stratospheric and tropospheric aircraft pollution on stratospheric ozone are now being investigated (section 9.2.2.2). The research studies have reemphasized that our understanding of stratospheric transport, is not complete with regard to transport, especially the containment of the pollutants within the midlatitude release regions. Many of the issues associated with the stratospheric circulation (section 9.2.1.1) are above the observing range of current stratospheric aircraft (i.e. above 70 mb). The analysis of UARS measurements has also revealed the tremendous advantages of global chemical data sets.
Finally, the most extensive observations of solar UV and energetic particle impact on ozone have been made recently by UARS. Unfortunately, these observations have been made during the declining phase of the solar cycle, and we have not developed a long enough baseline of measurements to quantify the impacts of changing solar conditions. Long term measurements of solar UV and total solar irradiance are needed during the waxing phase of the solar cycle.
The measurement requirements are discussed below. Table 9.3-1 summarizes the minimum measurements, their accuracies and the instruments which will make the measurements. Often, key measurements will be made by more than one instrument which gives the whole measurement suite a level of robustness in case of instrument failure.
An understanding of the photochemistry of the stratosphere is clearly contingent on high quality observations of temperature. The temperature field affects stratospheric physical processes in a number of ways. First, temperature fields are used to calculate geopotential heights and winds via the hydrostatic and geostrophic approximations. Second, temperatures affect the radiation field, particularly in relation to the longwave cooling in the stratosphere. Third, temperatures affect the chemistry via temperature dependent reaction rates, and via the formation of polar stratospheric clouds (the indirect cause of the ozone hole). Hence, accurate and precise temperatures provide a basic foundation for stratospheric chemistry, radiation, dynamics, and transport.
As stated in section 9.2.1.1 low quality tropical meteorological observations are an impediment to our understanding of the interaction of the tropics and the middle latitudes.
The National Plan for Stratospheric Monitoring 1988-1997 (1989) set down the minimum requirements for meteorological variables between 1000 and 0.1 hPa. Their requirements were: 2.7 km vertical resolution; 12 hour time resolution; 1 K precision for temperature, and 5 meter/sec precision for winds.
Current radiosonde and rawinsonde measurements have a few tenths of a degree K and 1-4 meter/sec wind speed in precision (Nash and Schmidlin, 1987). Unfortunately the balloon borne rawinsonde system is limited to altitudes below 30 km. For higher altitudes, the meteorological rocket network provided some data, but the network has been effectively discontinued. Satellite systems are now relied upon to provide all of the meteorological information above 30 km.
The NOAA TOVS (MSU, SSU and HIRS) SSU instrument has an error of 2K at 10 mb rising to 4K at 1 mb. The TOVS weighting functions are about 10-12 km deep. Future NOAA sounders will use AMSU instead of MSU/SSU. The AMSU weighting functions are about half the depth of the TOVS functions. AMSU will be providing data in 1996, but no AMSU temperature measurements are limited to the atmosphere below 50 mb..
UARS MLS limb sounder has vertical resolution of a few km although its horizontal coverage is inferior to nadir sounding TOVS and AMSU instruments. Improved understanding of stratospheric chemistry and heterogeneous processing suggests that improvement of temperature measurements will have an impact on our ability to predict where the heterogeneous reactions will take place.
Lower stratospheric temperature measurements made during the numerous polar aircraft missions suggests that the meteorological analysis (based upon TOVS) in the southern hemisphere are warm biased by about 2K. This suggests that the EOS stratospheric temperature accuracy requirements should be less than 0.5K. It is also important that good temperature measurements be made near the tropical tropopause especially in cloudy regions where air is entering the stratosphere through the tropopause.
Direct stratospheric wind data are needed where the divergence fields are significant ( e.g. the tropics). The current requirement for assimilation models is unbiased horizontal wind field accuracy of 2-5 m/s. These should be global measurements with a vertical resolution of a few kilometers. Presently, the UARS HRDI satellite wind instrument makes these measurements at the upper end of the limit.
Atmospheric composition measurements form a cornerstone of any global change strategy. Chemical and dynamical measurements must be made in both the stratosphere and the troposphere. Indeed, chemical measurements around the upper troposphere and lower stratosphere should be among those with the highest priority.
The science questions for stratospheric processes are mostly focused on the changes in the stratosphere expected to take place as anthropogenic pollutants accumulate in the middle atmosphere. Greenhouse gases are expected to substantially increase during the EOS period. Stratospheric halogens are expected to increase until 1999 then level off and slowly decline as a result of international regulations. The increases in these gases should produce chemical and dynamical changes. The magnitude of the stratospheric cooling in response to increasing greenhouse gases should far exceed the tropospheric warming because there are fewer feedback mechanisms (such as clouds) which buffer the radiative interaction. Increases in chlorine and bromine will cause decreases in stratospheric ozone. The ozone decrease could be exacerbated by colder lower stratospheric temperatures caused by increasing greenhouse gas concentrations. For example, the colder stratospheric temperatures may lead to an expansion of the extent of PSC's and, hence, polar ozone depletion.
The complex chemistry of the stratosphere can only be understood in detail by measuring a broad range of species over varying conditions with global coverage and over at least an annual cycle. The first area which merits further observational and theoretical study is polar chemistry processes. Direct, simultaneous measurements of HOCl (or a proxy such as ClO), HNO3, and N2O5 are critical since these gases are believed to be involved in PSC surface chemistry. Also, polar night observations, above 20 km, of the chemically active species, along with PSC measurements, are needed in understanding polar ozone depletion. These regions are not presently accessible with balloons and aircraft.
In order to understand the large ozone depletion at midlatitudes (see section 9.1.2.2), simultaneous measurements of N2O5, HOCl, HNO3, and HCl are needed to assess the role of heterogeneous chemistry on background aerosols. Since OH and HO2 drive the chemistry of the lower stratosphere, global measurements of these gases are required to evaluate ozone losses, especially any zonal asymmetries. Also, lower mesosphere observations of OH and HO2, along with O3 and temperature, are likely to be key links in understanding the large O3 decrease expected to occur near 40 km as chlorine levels continue to rise. It is clear that full understanding of these changes requires not just O3 and ClO measurements, but HOx and NOx as well. Measurements in the lower mesosphere where, the chemistry is more simple, may provide the best data set for this analysis.
There are several scientific requirements to address middle atmosphere chemistry issues:
A. The self-consistency in the source gases and the resulting active reservoir gases needs to be tested for the four major families that are important to ozone chemistry. The four families and most important species' measurements required are: oxygen family (O3), hydrogen family (H2O, CH4, OH, HO2, H2O2), nitrogen family (N2O, NO2, HNO3, N2O5), and chlorine family (CFCl3, CF2Cl2, HCl, ClO, ClONO2). Stratospheric chlorine is predicted by atmospheric models to increase by 20% in the next five years thus our understanding of the production and partitioning among the individual family constituents needs to be verified.
B. The changes in the Antarctic/Arctic lower stratosphere constituents (O3, H2O, ClO, OClO,HCl, BrO, N2O, NO2, HNO3, N2O5, and aerosols) during the ozone hole period in the winter and spring need to be monitored. Since significant changes have been detected during the 1980's and 90's in the polar regions, these geographical areas require special attention and monitoring.
C. There are a few chemical process studies which require investigation as indicated below.
C1. The HOx family (OH, HO2, H2O2) is fundamentally important in stratospheric chemistry, but the data base for that group remains one of the poorest in the atmosphere. Global measurements of the latitudinal, seasonal, and diurnal variation in the HOx family and related species, H2O and O3, are needed to address this deficiency.
C2. Models for the past decade have predicted less ozone in the upper stratosphere than is measured. Several species (O3, O, NO2, OH, and ClO) need to be measured in the upper stratosphere to help resolve this difficulty (see section 9.1.4.1).
C3. Models, in general, predict less odd nitrogen in the lower stratosphere than observed. Measurements of odd nitrogen species NO2, HNO3, N2O5, and ClONO2 in the lower stratosphere will help to deal with this problem.
C4. Another odd nitrogen species, HNO3, is not modeled accurately in the wintertime in the mid to high latitudes. A measurement of HNO3, N2O5, H2O, and aerosols should help confront this problem.
The measurement requirements to attack these science questions are outlined in Table 9.3-1. Generally, the accuracies needed are 5-10% of the ambient concentrations found in the lower stratosphere.
The remote sensing of the composition of aerosol at mid-latitudes is fairly straight forward. However, remote sensing of the composition and phase of the aerosol particles, for the case of the PSCs, is a developing topic of research. One research goal is to see to what extent it is possible to estimate the composition, phase, area and volume densities, from orbital observations. It is known that the volume densities of NAT and ternary particles are different. Carslaw et al. (l994) and Drdla et al. (1994) have shown that ternary solutions best describe some of the ER-2 data. For example, Figure 9.3.3.1 shows a graph of temperature versus volume density for several aerosol compositions. Beginning with a sulfuric acid droplet core, the volume density of the aerosol increases as temperatures become colder. These equilibrium curves were calculated using different amounts of ambient HNO3 (5, 10, and 15 ppbv). Remote sensing observations of temperature versus volume density (and/or aerosol extinction) will likely help classify the composition and phase of the PSC particles. It is also known that HNO3 is incorporated in ternary, NAD (nitric acid dihydrate), and NAT particles as a function of temperature (i.e., curves of temperature versus equilibrium gas phase HNO3 differ for the three compounds). Therefore, the simultaneous observation of aerosol extinction and HNO3 gas mixing ratios should help one to classify regions of PSCs as to composition and phase.
Since the microphysics of PSC particles is very temperature sensitive, absolute temperatures need to be measured to plus or minus 2K, since curves of temperature versus volume density for NAT, ternary, and NAD particles (Figure 9.3.3.1-1) differ by only a few degrees K. Remote sensing observations also average over many kilometers along a horizontal ray path. Vertical coverage is usually on the order of several km. Thus, the fine scale structure of PSCs, as sampled by ER-2instruments, can not be resolved by the remote sounder. Another complication is due to present limitations in the theoretical understanding of how PSCs form, which compositions are formed, and the need for additional laboratory work to quantify at cold stratospheric temperatures the rates at which realistic PSC particles convert inactive to active chlorine compounds, and the need for additional laboratory measurements of the refractive indices of PSC and sulfuric droplets. Current research will see to what extent it is possible to refine present capability to quantify the mechanisms of PSC chemistry, as observed from orbit.
Solar radiation at wavelengths below about 300 nm is completely absorbed by the earth's atmosphere and becomes the dominant direct energy input, establishing the composition and temperature through photodissociation, and driving much of the dynamics as well. Even small changes in this ultraviolet irradiance will have important and demonstrable effects on atmospheric ozone. Radiation between roughly 200 and 300 nm is absorbed by ozone and becomes the major loss mechanism for ozone in the middle atmosphere. Likewise solar radiation < 200 nm is absorbed predominantly by molecular oxygen and becomes a dominant source of ozone in the middle atmosphere so changes in these ultraviolet wavelengths will have, to first order, an inverse influence on ozone. These two atmospheric processes, driven by solar radiation, become the major natural control for ozone in the earth's stratosphere and lower thermosphere. To fully understand the ozone distribution will require many coordinated observations, and in particular, a precise measurement of the solar ultraviolet flux.
The visible portion of solar radiation originates in the solar photosphere and has been accurately measured for about fifteen years (Willson and Hudson, 1991). Apparently this radiation varies by only small fractions of one percent over the 11-year activity cycle of the Sun, with comparable variation over time scales of a few days. The ultraviolet portion of the solar spectrum comprises only about 1% ( approximately 10 W/m^2) and originates from higher layers of the photosphere. As we move to shorter and shorter wavelengths, the emission comes from higher and higher layers of the solar atmosphere. Unlike the solar photosphere, these higher levels are much more under the influence of solar activity, as manifested, for example, by increasing magnetic field strength. As the magnetic activity increases or disappears, the solar radiation, especially the ultraviolet, undergoes dramatic variations modulated by the 27-day rotation period of the Sun. Near 120 nm the variation over time periods of days to weeks can be as large as 50%, and over the longer 11-year solar cycle the variation can be as large as a factor of two (Rottman, 1988). Toward longer wavelengths, the solar variability decreases to levels of about 10% near at 200 nm (Figure 9.3.4-1) and finally to only about 1% at 300 nm. Longward of 300 nm, the intrinsic solar variability is probably only on the order 0.1%, roughly commensurate with measurements of total solar radiation.
The challenge during the EOS time period is to provide measurements of the solar ultraviolet with a precision and accuracy capable of tracking the changes in the solar output. Ideally the instrument will be capable of measuring changes as small as one percent throughout the EOS mission. This requirement is extremely challenging for solar instruments, especially those making observations at the ultraviolet wavelengths which are notoriously variable. The harsh environment of space, coupled with the energetic solar radiation, rapidly degrade optical surfaces and usually make the observations suspect. Some manner of in-flight calibration is required to unambiguously separate changes in the instrument response from true solar changes.
The role of validation of satellites-based chemical measurements can not be over stressed. Validation measurements, especially measurements of the same species using two different techniques, have proved to be invaluable for understanding satellite trace species measurements. The very successful UARS validation campaign has contributed a great deal to understanding of the individual UARS measurements. The validation campaigns perform two major functions. First, they test the ability of a satellite instrument to make a measurement by giving an independent data point to compare against. Second, if the validation measurements are performed as part of a larger, coordinated campaign, the validation measurements done using aircraft and ground-based measurements can be used to link the small scale geophysical features that they can observe with the large-scale geophysical features observable from space.
The tropopause, the boundary between the upper troposphere (UT) and lower stratosphere (LS), is critical for understanding many important processes in the atmosphere. The tropopause is defined by a sharp change in the vertical temperature gradient, taking place over a few hundred meters at most. Below the tropopause, the troposphere is a region of active vertical mixing. Above the tropopause, the stratosphere is very stable with little vertical mixing. The match between these dissimilar regions, troposphere and stratosphere, modulates the processes that permit the exchange of mass, trace gases, momentum, potential vorticity and energy between the two regions.
Unfortunately, present observing systems do not observe the UT-LS region with sufficient detail. The NOAA operational temperature sensors are characterized by vertical resolution of the retrievals are of the order of 10-12 km (Smith et al., 1979). The detailed structure of the tropopause is much too thin to be seen by operational systems. However, their cross-track scanning capability gives them the ability to observe horizontal scales of about 100 km (Figure 9.4.1-1).
Temperature profiles with much higher vertical resolution can be obtained by observing the atmospheric limb, or horizon (Gille and House, 1971). The improvement results from the geometry, since most of the ray path through the atmosphere is within 1-2 km of the lowest, or tangent, point. In addition, the atmospheric signal is seen against the cold background of space. These can reduce the height of the vertical weighting functions to 3-4 km (Bailey and Gille, 1986), and the effective resolution to ~ 5 km.
EOS limb sounders (MLS and HIRDLS on EOS CHEM-1) will greatly improve the accuracy, precision and resolution of temperature measurements in the tropopause region. HIRDLS will determine temperatures with a resolution of 1-1.5 km, through a combination of a narrow (1 km) vertical FOV, low noise, and oversampling. MLS will make limb temperature measurements with a resolution of 2-3 km.
Previous limb scanners have retrieved temperatures with higher vertical resolution, but, because the vertical scans are made at a single azimuth relative to the orbital plane. Thus the horizontal resolution is limited to the orbital spacing, or about 25 degrees, sufficient to resolve only about 6 longitudinal waves. However, the UTLS is a region in which smaller scale waves from thetroposphere are present. The horizontal resolution of previous limb sounders did not allow these smaller scale systems to be measured. Figure 9.4.1-1 shows that the consistent scaling between vertical and horizontal scales suggests that higher horizontal resolution is required to sample geostrophic motions. The figure shows that HIRDLS has the ability to observe and retrieve with a 1 km vertical resolution and 4o horizontal resolution by scanning from side to side. This resolution allows all horizontal waves, up to ~ wavenumber 45, to be observed with the appropriate vertical resolution.
Recently a technological innovation has been proposed for the MLS instrument. Instead of a single heterodyne receiver, MMIC arrays have been proposed at two frequencies. The array system (AMLS - Array MLS) would allow 100km by 100km horizontal resolution temperature, ozone, N2O, and water for the same cost, and lower power and weight. This proposed system is currently being studied by NASA.
EOS instruments will give significantly improved stratospheric chemical measurements by having better measurement precision, particularly in the lower stratosphere, and a more complete suite of collocated measurements, especially chemical radicals. The accuracy of the proposed instruments is close to that set down in Table 9.3-1. The major improvements in these chemical measurements are from MLS and HIRDLS. These two instruments (especially AMLS) are very synergistic in that HIRDLS will have high resolution in longitude as well as latitude (discussed in the previous section) while MLS will be able to make measurements in high aerosol or cloudy regions. In addition, high-latitude coverage will be obtained on each orbit from both HIRDLS and MLS, a significant improvement from UARS, which had monthly gaps in high latitude coverage and important periods were not sampled.
The high vertical and horizontal resolution coverage in the upper troposphere and lower stratosphere by HIRDLS (and AMLS) are extremely important because atmospheric waves with smaller horizontal scales can penetrate to these altitudes, creating variations on small scales that are critical to our understanding of wave breaking, mixing, and perhaps chemical processing.
Compared with UARS, EOS MLS has tremendous improvement in precision of lower stratospheric measurements due to its increased spectral bandwidth and choice of stronger spectral lines. Whereas UARS MLS was designed primarily for the middle and upper stratosphere, EOS MLS emphasizes the lower stratosphere. Improvements over UARS are summarized in TABLE 9.4.2.1-1 (IMPROVEMENTS IN EOSCHEM MEASUREMENTS OVER UARS)
Inclusion of the OH measurement in MLS is a major qualitative improvement in the suite of global stratospheric measurements which EOS will provide. This measurement is possible because of recent submillimeter-wavelength technology advances which were not available for UARS. The OH measurement will extend to the lowest and the highest regions of the stratosphere, where HOx chemistry is thought to be the dominant ozone loss mechanism on a global scale. It will cover regions where OH is thought to control the conversion of CH4 to H2O, to control the rates of SO2 and OCS oxidation to sulfate aerosol, and to be an essential player in controlling the balance between radical and reservoir species in the nitrogen and chlorine families. The global OH measurement over the complete stratosphere by EOS will give unprecedented new information on stratospheric chemistry, and is especially valuable in being made simultaneously with that of the many other EOS chemistry measurements. MLS will also likely be able to measure HO2 in the upper stratosphere, further testing and improving our understanding of stratospheric hydrogenchemistry.
HIRDLS NO2 measurements provide and important component of NOx, and, because of its reactions with ClO (measured by MLS) the formation of ClONO2, also measured by HIRDLS. HIRDLS measurements of N2O5 and HNO3, along with the ClONO2, provide a fairly complete set of measurements of the NOy species. Note that the ratio of NO2/HNO3 provides an additional way of deriving the OH concentration.
MLS will also measure middle and upper stratospheric BrO, the dominant radical in the bromine chemical family. No global stratospheric bromine measurements have been made to date, and the BrO measurement will be important to test our understanding of this chemistry.
Other important improvements in the suite of chemical measurements include the simultaneous and commonly- calibrated measurement of HCl and ClO by MLS. This allows very accurate monitoring of the ClO/HCl ratio which provides a sensitive indicator of our understanding of chlorine chemistry and early detection of changes. The MLS and HIRDLS N2O measurements will allow much more accurate distinction between chemical and dynamical changes in the atmosphere.
EOS instruments will improve upon the ability demonstrated by UARS for several reasons. Temperature retrievals will be more refined, and the horizontal and vertical resolution will be better. HIRDLS resolution will be on the order of 1 km in the vertical coordinate, and 4 degrees latitude by 4 degrees longitude in the horizontal coordinate, and will retrieve temperatures with an accuracy of 1K at altitudes below 50 km (Gille and Barnett, 1992). The HIRDLS experiment has four spectral channels which are specifically dedicated to obtaining aerosol extinction measurements, but will obtain aerosol information from many of the other channels as well. The measurements will observe and distinguish mid-latitude sulfuric acid droplets and PSCs. H2O and HNO3, two gases which are incorporated into stratospheric aerosols, will be retrieved by HIRDLS. In addition, ClONO2 and N2O5, species important in heterogeneous processing, are retrieved by HIRDLS.
The SAGE III aerosol data will consist of vertical profiles of aerosol extinction at seven wavelengths extending in altitude range from the lower troposphere or cloud top to about 40 KM altitude, with vertical resolution of one km. From the seven wavelength aerosol extinction measurements, parameters describing the aerosol physical size distribution such as volume density and surface area density, can be estimated with good accuracy, with uncertainties of the order of 10%. The wide spectral extinction behavior from the SAGE III measurements should also provide sufficient information to distinguish between a single-modal versus a bimodal aerosol size distribution, especially important immediately after volcanic eruptions. SAGE III instrument will provide better tropospheric aerosol measurements with extinction measurements with multiple wavelengths through the mid-troposphere, and at least at two wavelengths down to sea level. For the measurements of Polar Stratospheric Clouds (PSCs), SAGE III will be the first satellite instrument to provide size distribution information.
The SAGE III aerosol information will be key to understanding the role played by stratospheric aerosols and PSCs in heterogeneous chemical reactions and ozone depletion and will be indispensable for understanding aerosol radiative forcing.
The EOS SOLSTICE has the unique capability of observing bright, blue stars employing the very same optics and detectors used for the solar observations. With the assumption that these stars are extremely stable, the response of the SOLSTICE could be accurately monitored by observing a single star. However the calibration plan uses twenty or more of these reference stars, and it is the average flux from the entire ensemble of stars that provides an even more reliable reference for the solar observations. Repeated observation of the stars should yield an average flux level that is unchanging in time. A decrease in the level is an indication of loss in instrument sensitivity and adjustments are made accordingly. Since the same optics and detectors are used for the solar observations, the same corrections apply and the resulting solar measurement is reliable and free of any instrumental effects. The large dynamic range between the stellar and solar signals is on the order of eight to nine orders of magnitude, and is easily accommodated by changing the entrance apertures, spectral bandpass and integration times (Rottman et al., 1993).
EOS is currently funding two efforts and cooperates with a foreign effort to build fully interactive chemical/dynamical models for interpretation of EOS data. The first effort, the Schoeberl IDS, currently uses data assimilation winds and temperatures (see section 9.4.6) along a chemical transport model to simulate the stratospheric system. The results of this model are then used to compare with observations such as those from UARS, SAGE and TOMS data. Full coupling and feedback between the meteorology and chemistry does not occur at this stage. Later, the chemical/transport model will be used along with the EOS chemical data to produce an assimilated chemical data set. The full coupling will occur through the assimilation system. The observations of motion of long lived trace species can be inverted to make better estimates of stratospheric winds. The forecast of temperatures from the assimilation modeling effort will increase the accuracy of the retrieved trace gases.
The second effort, the Pyle IDS (United Kingdom), uses a full GCM and chemical package. The GCM and chemical/radiation/transport system will be fully linked. The objectives are to examine the sensitivity of the atmosphere to the chemical/radiative/dynamical feedback systems. This foreign effort is comprised of a full fledged EOS IDS team but is not funded by NASA.
The third effort, the Grose IDS, is also be developing a fully coupled general circulation model including a comprehensive formulation for the chemistry of the stratosphere and troposphere. An off-line CTM is currently used in conjunction with a variable resolution general circulation model. Multi-year simulations have been performed at several different horizontal and vertical resolutions with this version of the model. The model has been used in comparisons with UARS data (Eckman et al., 1995), and studies of stratospheric dynamics and transport (Pierce, et al., 1993). In a parallel effort, progress has been made towards the goal of a fully coupled model to address the interactions between ozone radiative heating and transport. In recent simulations, transported ozone is used in the stratospheric longwave and shortwave heating calculations.
The Data Assimilation Office (DAO) at NASA/GSFC already provides routine support of the stratospheric aircraft missions planned to study stratospheric ozone. Daily analyses are produced for the STRAT mission, and given the duration of this mission, this will evolve into the operational support of the AM-1 Platform. During the aircraft deployment periods, forecasts are produced to aid in flight planning. These forecasts enhance the ability of the mission planners to target air of specific chemical characteristics.
The current DAO meteorological analyses have a large impact on stratospheric chemistry studies. The wind fields are of sufficient quality to remove the dynamical uncertainty from tracer observations. This allows a high caliber examination of chemical processes, a benefit that has been realized for satellite, balloon, and aircraft data. There are active efforts of the DAO to improve the subtropical winds and the deep vertical motions that link the stratosphere and mesosphere.
Future plans call for a straightforward extension of the application of winds from the DAO assimilation to more general problems. This includes stratospheric-tropospheric exchange, as well as broader issues of tropospheric chemistry. Advanced assimilation systems will directly assimilate of constituents observations. First, long-lived tracers will be assimilated. Initial studies of N2O from UARS show that assimilation can in fact provide verifiable global information from the non-global UARS coverage pattern. An important goal of the DAO constituent effort is to improve wind estimates, especially in the tropics. Wind inversion techniques are being actively investigated, and we expect to have the first results in 1997. In addition, at least two university proposals have been recently submitted to attempt to assimilate aerosol observations within the DAO system. These proposals are examples of the long-term efforts to assimilate the complete chemical suite of measurements, with the goal of bringing to bear the quantitative analysis of data assimilation on the internal consistency of the chemical observations of UARS and CHEM-1.
Foreign partners are critical to the scientific success of the global stratospheric measurements program. The scale of collaboration ranges from individual science teams to collaborative instruments to reciprocal flights of instruments to mission planning. The timing of the EOS CHEM mission has been structured to follow the launch of ESA's ENVISAT I. This mission, to be launched in mid-1998, follows the successful NASA UARS mission. ENVISAT will make many critical trace species measurements in the time period of maximum stratospheric chlorine. This long-term data set, UARS - ENVISAT - CHEM, will be absolutely critical to our understanding of the role of trace species in controlling ozone in the stratosphere.
Other international space platforms will carry a few stratospheric instruments: The Japanese ADEOS platform will carry the Total Ozone Mapping Spectrometer (TOMS) and the French SPOT satellite will carry POAM II aerosol and ozone measuring system. The Russian MUIR space station will fly the Fourier Transform Spectrometer, DOPI.
In addition to these space instruments, several stratospheric aircraft campaigns are planned for the next few years (e.g. the Stratospheric Tracers of Atmospheric Transport (STRAT) campaign).
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Figure9.1.1-1 The distribution of atmospheric ozone in partial pressure as a function of altitude.
Figure 9.1.1-2 The global distribution of column or total ozone averaged over the 13 years of Nimbus-7 TOMS data in Dobson Units (DU).
Figure 9.1.2-1 An update of the Halley Bay column ozone data for October with TOMS and BUV satellite measurements.
Figure 9.1.2-2 Balloonsonde ozone measurements made at McMurdo and SAGE measurements for 1987.
Figure 9.1.2-3 Time series of column ozone measurements at Arosa, Switzerland in DU.
Figure 9.1.3.3.2-1 SAM II measurements of vertical optical depth data in the stratosphere over the Arctic and Antarctic. The measurements show the impact of volcanic eruptions over the period 1979-1992 along with seasonal effects in the local winters due to Polar Stratospheric Clouds (PSCs). The data are weekly averages at a wavelength of 1000 nm.
Figure 9.2.2.1.2-1 Trend in HALOE HCl vs. time at 55km compared with the estimated trend based on tropospheric emissions.
Figure 9.2.2.1.3-1 A summary of model calculations of the percent change in column ozone verses time for March at 50N for Scenario 1(reproduced from Fig. 6-12 of WMO, 1995).
Figure 9.3.3-1 Temperature versus volume density for several aerosol compositions expected in the stratosphere.
Figure 9.3.4-1 Changes in solar ultraviolet flux verses time from the UARS SOLSTICE instrument.
Figure 9.4.1-1 A comparison of the vertical and horizontal of various limb and nadir sounding satellite instruments. Lines indicate the ideal observation scale for Rossby waves.
Figure 9.4.3-1 Part a shows UKMO temperatures for January 10, 1992 at 46 mb. Aerosol data for from UARS CLAES (Part b) at the same level for the same date. Note the appearance of aerosols in the coldest region.
STRATOSPHERIC CHEMICAL AND
DYNAMICAL MEASUREMENT REQUIREMENTS
Measurement Accuracy EOS Instrument Meteorology Temperature 1K MLS, HIRDLS Winds 2-5 m/s (none) Chemistry O3 0.2 ppm MLS, HIRDLS, SAGE III H2O 0.5 ppm MLS, HIRDLS, SAGE III CFC-11/12 0.2 ppb HIRDLS N2O 20 ppb MLS, HIRDLS CH4 0.1 ppm HIRDLS HCl 0.1 ppb MLS ClONO2 0.1 ppb HIRDLS HNO3 1.0 ppb HIRDLS, MLS NO2/NO 0.2 ppb HIRDLS ClO 50 ppt MLS BrO 5 ppt MLS OH 0.5 ppt MLS N2O5 0.2 ppb HIRDLS Aerosols Surface area SAGE III, HIRDLS within 10% Solar Flux 100-400 nm SOLSTICE to 4% Requirements include vertical resolution of 1-2 km through the tropopause into the lower stratosphere. Horizontal resolution is minimally that of UARS (2700 km) but increased horizontal resolution vastly improves the science. HIRDLS scanning will achieve a horizontal resolution of 500 km, AMLS may achieve a horizontal resolution of 100 km.
