Ozone Layer Solid Research Subject: TABLE OF CONTENTS How to get this FAQ Copyright Statement General remarks Caveats, Disclaimers, and Contact Information TABLE OF CONTENTS 1. THE STRATOSPHERE 1.1) What is the stratosphere? 1.2) How is the composition of air described? 1.3) How does the composition of the atmosphere change with 2. THE OZONE LAYER 2.1) How is ozone created? 2.2) How much ozone is in the layer, and what is a 2.3) How is ozone distributed in the stratosphere? 2.4) How does the ozone layer work? 2.5) What sorts of natural variations does the ozone layer show? 2.5.a) Regional and Seasonal Variation 2.5.b) Year-to-year variations. 2.6) What are CFC’s? 2.7) How do CFC’s destroy ozone? 2.8) What is an Ozone Depletion Potential? 2.9) What about HCFC’s and HFC’s? Do they destroy ozone? 2.10) *IS* the ozone layer getting thinner? 2.11) Is the middle-latitude ozone loss due to CFC emissions? 2.12) If the ozone is lost, won’t the UV light just penetrate 2.13) Do Space Shuttle launches damage the ozone layer? 2.14) Will commercial supersonic aircraft damage the ozone layer? 2.15) What is being done about ozone depletion? 3. REFERENCES FOR PART I Introductory Reading Books and Review Articles More Specialized References Internet Resources —————————– Subject: 1. THE STRATOSPHERE —————————– Subject: 1.1) What is the stratosphere? The stratosphere extends from about 15 km to 50 km.
In the stratosphere temperature increases with altitude, due to the absorption of UV light by oxygen and ozone. This creates a global inversion layer which impedes vertical motion into and within the stratosphere – since warmer air lies above colder air, convection is inhibited. The word stratosphere is related to the word stratification or layering. The stratosphere is often compared to the troposphere, which is the atmosphere below about 15 km. The boundary – called the tropopause – between these regions is quite sharp, but its precise location varies between ~9 and ~18 km, depending upon latitude and season.
The prefix tropo refers to change: the troposphere is the part of the atmosphere in which weather occurs. This results in rapid mixing of tropospheric air. [Wayne] [Wallace and Hobbs] Above the stratosphere lie the mesosphere, ranging from ~50 to ~100 km, in which temperature decreases with altitude; the thermosphere, ~100-400 km, in which temperature increases with altitude again, and the exosphere, beyond ~400 km, which fades into the background of interplanetary space. In the upper mesosphere and thermosphere electrons and ions are abundant, so these regions are also referred to as the ionosphere. In technical literature the term lower atmosphere is synonymous with the troposphere, middle atmosphere refers to the stratosphere and mesosphere, while upper atmosphere is usually reserved for the thermosphere and exosphere.
This usage is not universal, however, and one occasionally sees the term upper atmosphere used to describe everything above the troposphere (for example, in NASA’s Upper Atmosphere Research Satellite, UARS.) —————————– Subject: 1.2) How is the composition of air described? (Or, what is a ‘mixing ratio’?) The density of the air in the atmosphere depends upon altitude, and in a complicated way because the temperature also varies with altitude. It is therefore awkward to report concentrations of atmospheric species in units like g/cc or molecules/cc. Instead, it is convenient to report the mole fraction, the relative number of molecules of a given type in an air sample. Atmospheric scientists usually call a mole fraction a mixing ratio. Typical units for mixing ratios are parts-per-million, billion, or trillion by volume, designated as ppmv, ppbv, and pptv respectively. (The expression by volume reflects Avogadro’s Law – for an ideal gas mixture, equal volumes contain equal numbers of molecules – and serves to distinguish mixing ratios from mass fractions which are given as parts-per-million by weight.) Thus when someone says the mixing ratio of hydrogen chloride at 3 km is 0.1 ppbv, he means that 1 out of every 10 billion molecules in an air sample collected at that altitude will be an HCl molecule.
[Wayne] [Graedel and Crutzen] —————————– Subject: 1.3) How does the composition of the atmosphere change with altitude? (Or, how can CFC’s get up to the stratosphere when they are heavier than air?) In the earth’s troposphere and stratosphere, most stable chemical species are well-mixed – their mixing ratios are independent of altitude. If a species’ mixing ratio changes with altitude, some kind of physical or chemical transformation is taking place. That last statement may seem surprising – one might expect the heavier molecules to dominate at lower altitudes. The mixing ratio of Krypton (mass 84), then, would decrease with altitude, while that of Helium (mass 4) would increase. In reality, however, molecules do not segregate by weight in the troposphere or stratosphere.
The relative proportions of Helium, Nitrogen, and Krypton are unchanged up to about 100 km. Why is this? Vertical transport in the troposphere takes place by convection and turbulent mixing. In the stratosphere and in the mesosphere, it takes place by eddy diffusion – the gradual mechanical mixing of gas by motions on small scales. These mechanisms do not distinguish molecular masses. Only at much higher altitudes do mean free paths become so large that molecular diffusion dominates and gravity is able to separate the different species, bringing hydrogen and helium atoms to the top.
The lower and middle atmosphere are thus said to be well mixed. [Chamberlain and Hunten] [Wayne] [Wallace and Hobbs] Experimental measurements of the fluorocarbon CF4 demonstrate this homogeneous mixing. CF4 has an extremely long lifetime in the stratosphere – probably many thousands of years. The mixing ratio of CF4 in the stratosphere was found to be 0.056-0.060 ppbv from 10-50 km, with no overall trend. [Zander et al. 1992] An important trace gas that is *not* well-mixed is water vapor. The lower troposphere contains a great deal of water – as much as 30,000 ppmv in humid tropical latitudes.
High in the troposphere, however, the water condenses and falls to the earth as rain or snow, so that the stratosphere is extremely dry, typical mixing ratios being about 5 ppmv. Indeed, the transport of water vapor from troposphere to stratosphere is even less efficient than this would suggest, since much of the small amount of water in the stratosphere is actually produced in situ by the oxidation of stratospheric methane. [SAGE II] Sometimes that part of the atmosphere in which the chemical composition of stable species does not change with altitude is called the homosphere. The homosphere includes the troposphere, stratosphere, and mesosphere. The upper regions of the atmosphere – the thermosphere and the exosphere – are then referred to as the heterosphere.
[Wayne] [Wallace and Hobbs] —————————– Subject: 2. THE OZONE LAYER —————————– Subject: 2.1) How is ozone created? Ozone is formed naturally in the upper stratosphere by short wavelength ultraviolet radiation. Wavelengths less than ~240 nanometers are absorbed by oxygen molecules (O2), which dissociate to give O atoms. The O atoms combine with other oxygen molecules to make ozone: O2 + hv -* O + O (wavelength * 240 nm) O + O2 -* O3 —————————– Subject: 2.2) How much ozone is in the layer, and what is a Dobson Unit ? A Dobson Unit (DU) is a convenient scale for measuring the total amount of ozone occupying a column overhead. If the ozone layer over the US were compressed to 0 degrees Celsius and 1 atmosphere pressure, it would be about 3 mm thick. So, 0.01 mm thickness at 0 C and 1 at is defined to be 1 DU; this makes the average thickness of the ozone layer over the US come out to be about 300 DU. In absolute terms, 1 DU is about 2.7 x 10^16 molecules/cm^2.
The unit is named after G.M.B. Dobson, who carried out pioneering studies of atmospheric ozone between ~1920-1960. Dobson designed the standard instrument used to measure ozone from the ground. The Dobson spectrophotometer measures the intensity solar UV radiation at four wavelengths, two of which are absorbed by ozone and two of which are not [Dobson 1968b]. These instruments are still in use in many places, although they are gradually being replaced by the more elaborate Brewer spectrophotometers. Today ozone is measured in many ways, from aircraft, balloons, satellites, and space shuttle missions, but the worldwide Dobson network is the only source of long-term data.
A station at Arosa in Switzerland has been measuring ozone since the 1920’s (see http://www.umnw.ethz.ch/LAPETH/doc/totozon.html) and some other stations have records that go back nearly as long, although many were interrupted during World War II. The present worldwide network went into operation in 1956-57. —————————– Subject: 2.3) How is ozone distributed in the stratosphere? In absolute terms: about 10^12 molecules/cm^3 at 15 km, rising to nearly 10^13 at 25 km, then falling to 10^11 at 45 km. In relative terms: ~0.5 parts per million by volume (ppmv) at 15 km, rising to ~8 ppmv at ~35 km, falling to ~3 ppmv at 45 km. Even in the thickest part of the layer, ozone is a trace gas. In all, there are about 3 billion metric tons, or 3×10^15 grams, of ozone in the earth’s atmosphere; about 90% of this is in the stratosphere. —————————– Subject: 2.4) How does the ozone layer work? UV light with wavelengths between 240 and 320 nm is absorbed by ozone, which then falls apart to give an O atom and an O2 molecule. The O atom soon encounters another O2 molecule, however (at all times, the concentration of O2 far exceeds that of O3), and recreates O3: O3 + hv -* O2 + O O + O2 -* O3 Thus ozone absorbs UV radiation without itself being consumed ; the net result is to convert UV light into heat.
Indeed, this is what causes the temperature of the stratosphere to increase with altitude, giving rise to the inversion layer that traps molecules in the troposphere. The ozone layer isn’t just in the stratosphere; the ozone layer actually determines the form of the stratosphere. Ozone is destroyed if an O atom and an O3 molecule meet: O + O3 -* 2 O2 (recombination). This reaction is slow, however, and if it were the only mechanism for ozone loss, the ozone layer would be about twice as thick as it is. Certain trace species, such as the oxides of Nitrogen (NO and NO2), Hydrogen (H, OH, and HO2) and chlorine (Cl, ClO and ClO2) can catalyze the recombination.
The present ozone layer is a result of a competition between photolysis and recombination; increasing the recombination rate, by increasing the concentration of catalysts, results in a thinner ozone layer. Putting the pieces together, we have the set of reactions proposed in the 1930’s by Sidney Chapman: O2 + hv -* O + O (wavelength * 240 nm) : creation of oxygen atoms O + O2 -* O3 : formation of ozone O3 + hv -* O2 + O (wavelength * 320 nm) : absorption of UV by ozone O + O3 -* 2 O2 : recombination . Since the photolysis of O2 requires UV radiation while recombination does not, one might guess that ozone should increase during the day and decrease at night. This has led some people to suggest that the antarctic ozone hole is merely a result of the long antarctic winter nights. This inference is incorrect, because the recombination reaction requires oxygen atoms which are also produced by photolysis. Throughout the stratosphere the concentration of O atoms is orders of magnitude smaller than the concentration of O3 molecules, so both the production and the destruction of ozone by the above mechanisms shut down at night.
In fact, the thickness of the ozone layer varies very little from day to night, and above 70 km ozone concentrations actually increase at night. (The unusual catalytic cycles that operate in the antarctic ozone hole do not require O atoms; however, they still require light to operate because they also include photolytic steps. See Part III.) —————————– Subject: 2.5) What sorts of natural variations does the ozone layer show? There are substantial variations from place to place, and from season to season. There are smaller variations on time scales of years and more. [Wayne] [Rowland 1991] We discuss these in turn. —————————– Subject: 2.5.a) Regional and Seasonal Variation Since solar radiation makes ozone, one expects to see the thickness of the ozone layer vary during the year.
This is so, although the details do not depend simply upon the amount of solar radiation received at a given latitude and season – one must also take atmospheric motions into account. (Remember that both production and destruction of ozone require solar radiation.) The ozone layer is thinnest in the tropics, about 260 DU, almost independent of season. Away from the tropics seasonal variations become important. For example: Location Column thickness, Dobson Units Jan Apr Jul Oct Huancayo, Peru (12 degrees S) : 255 255 260 260 Aspendale, Australia (38 deg. S): 300 280 335 360 Arosa, Switzerland (47 deg. N): 335 375 320 280 St.
Petersburg, Russia (60 deg. N): 360 425 345 300 These are monthly averages. Interannual standard deviations amount to ~5 DU for Huancayo, 25 DU for St. Petersburg. [Rowland 1991]. Day-to-day fluctuations can be quite large (as much as 60 DU at high latitudes).
Notice that the highest ozone levels are found in the spring , not, as one might guess, in summer, and the lowest in the fall, not winter. Indeed, at high latitudes in the Northern Hemisphere there is more ozone in January than in July! Most of the ozone is created over the tropics, and then is carried to higher latitudes by prevailing winds (the general circulation of the stratosphere.) [Dobson 1968a] [Garcia] [Salby and Garcia] [Brasseur and Solomon] The antarctic ozone hole, discussed in detail in Part III, falls far outside this range of natural variation. Mean October ozone at Halley Bay on the Antarctic coast was 117 DU in 1993, down from 321 DU in 1956. —————————– Subject: 2.5.b) Year-to-year variations. Since ozone is created by solar UV radiation, one expects to see some correlation with the 11-year solar sunspot cycle. Higher sunspot activity corresponds to more solar UV and hence more rapid ozone production. This correlation has been verified, although its effect is small, about 2% from peak to trough averaged over the earth, about 4% in polar regions. [Stolarski et al.] Another natural cycle is connected with the quasibiennial oscillation, in which tropical winds in the lower stratosphere switch from easterly to westerly with a period of about two years.
This leads to variations of the order of 3% at a given latitude, although the effect tends to cancel when one averages over the entire globe. Episodes of unusual solar activity (solar proton events) can also influence ozone levels, by producing nitrogen oxides in the upper stratosphere and mesosphere. This can have a marked, though short-lived, effect on ozone concentrations at very high altitudes, but the effect on total column ozone is usually small since most of the ozone is found in the lower and middle stratosphere. Ozone can also be depleted by a major volcanic eruption, such as El Chichon in 1982 or Pinatubo in 1991. The principal mechanism for this is not injection of chlorine into the stratosphere, as discussed in Part II, but rather the injection of sulfate aerosols which change the radiation balance in the stratosphere by scattering light, and which convert inactive chlorine compounds to active, ozone-destroying forms.
[McCormick et al. 1995]. This too is a transient effect, lasting 2-3 years. —————————– Subject: 2.6) What are CFC’s? CFC’s – ChloroFluoroCarbons – are a class of volatile organic compounds that have been used as refrigerants, aerosol propellants, foam blowing agents, and as solvents in the electronic industry. They are chemically very unreactive, and hence safe to work with. In fact, they are so inert that the natural reagents that remove most atmospheric pollutants do not react with them, so after many years they drift up to the stratosphere where short-wave UV light dissociates them.
CFC’s were invented in 1928, but only came into large-scale production after ~1950. Since that year, the total amount of chlorine in the stratosphere has increased by a factor of 4. [Solomon] The most important CFC’s for ozone depletion are: Trichlorofluoromethane, CFCl3 (usually called CFC-11 or R-11); Dichlorodifluoromethane, CF2Cl2 (CFC-12 or R-12); and 1,1,2 Trichlorotrifluoroethane, CF2ClCFCl2 (CFC-113 or R-113). R stands for refrigerant. One occasionally sees CFC-12 referred to as F-12, and so forth; theF stands for Freon, DuPont’s trade name for these compounds. In discussing ozone depletion, CFC is occasionally used to describe a somewhat broader class of chlorine-containing organic compounds that have similar properties – unreactive in the troposphere, but readily photolyzed in the stratosphere.
These include: HydroChloroFluoroCarbons such as CHClF2 (HCFC-22, R-22); Carbon Tetrachloride (tetrachloromethane), CCl4; Methyl Chloroform (1,1,1 trichloroethane), CH3CCl3 (R-140a); and Methyl Chloride (chloromethane), CH3Cl. (The more careful publications always use phrases like CFC’s and related compounds, but this gets tedious.) Only methyl chloride has a large natural source; it is produced biologically in the oceans and chemically from biomass burning. The CFC’s and CCl4 are nearly inert in the troposphere, and have lifetimes of 50-200+ years. Their major sink is photolysis by UV radiation. [Rowland 1989, 1991] The hydrogen-containing halocarbons are more reactive, and are removed in the troposphere by reactions with OH radicals.
This process is slow, however, and they live long enough (1-20 years) for a substantia fraction to reach the stratosphere. Most of Part II is devoted to stratospheric chlorine chemistry; look there for more detail. —————————– Subject: 2.7) How do CFC’s destroy ozone? CFC’s themselves do not destroy ozone; certain of their decay products do. After CFC’s are photolyzed, most of the chlorine eventually ends up as Hydrogen Chloride, HCl, or Chlorine Nitrate, ClONO2. These are called reservoir species – they do not themselves react with ozone.
However, they do decompose to some extent, giving, among other things, a small amount of atomic chlorine, Cl, and Chlorine Monoxide, ClO, which can catalyze the destruction of ozone by a number of mechanisms. The simplest is: Cl + O3 -* ClO + O2 ClO + O -* Cl + O2 Net effect: O3 + O -* 2 O2 Note that the Cl atom is a catalyst – it is not consumed by the reaction. Each Cl atom introduced into the stratosphere can destroy thousands of ozone molecules before it is removed. The process is even more dramatic for Bromine – it has no stable reservoirs, so the Br atom is always available to destroy ozone. On a per-atom basis, Br is 10-100 times as destructive as Cl. On the other hand, chlorine and bromine concentrations in the stratosphere are very small in absolute terms.
The mixing ratio of chlorine from all sources in the stratosphere is about 3 parts per billion, (most of which is in the form of CFC’s that have not yet fully decomposed) whereas ozone mixing ratios are measured in parts per million. Bromine concentrations are about 100 times smaller still. (See Part II.) The complete chemistry is very complicated – more than 100 distinct species are involved. The rate of ozone destruction at any given time and place depends strongly upon how much Cl is present as Cl or ClO, and thus upon the rate at which Cl is released from its reservoirs. This makes quantitative predictions of future ozone depletion difficult. [Rowland 1989, 1991] [Wayne] The catalytic destruction of ozone by Cl-containing radicals was first suggested by Richard Stolarski and Ralph Cicerone in 1973. However, they were not aware of any large sources of stratospheric chlorine. In 1974 F.
Sherwood Rowland and Mario Molina realized that CFC’s provided such a source. [Molina and Rowland 1974][Rowland and Molina 1975] For this and for their many subsequent contributions to stratospheric ozone chemistry Rowland and Molina shared the 1995 Nobel Prize in Chemistry, together with Paul Crutzen, discoverer of the NOx cycle. (The official announcement from the Swedish Academy can be found on the web at http://www.nobel.se/announcement95-chemistry.html .) —————————– Subject: 2.8) What is an Ozone Depletion Potential? The ozone depletion potential (ODP) of a compound is a simple measure of its ability to destroy stratospheric ozone. It is a relative measure: the ODP of CFC-11 is defined to be 1.0, and the ODP’s of other compounds are calculated with respect to this reference point. Thus a compound with an ODP of 0.2 is, roughly speaking, one-fifth as bad as CFC-11. More precisely, the ODP of a compound x is defined as the ratio of the total amount of ozone destroyed by a fixed amount of compound x to the amount of ozone destroyed by the same mass of CFC-11: Global loss of Ozone due to x ODP(x) == ——————————— Global loss of ozone due to CFC-11.
Thus the ODP of CFC-11 is 1.0 by definition. The right-hand side of the equation is calculated by combining information from laboratory and field measurements with atmospheric chemistry and tranport models. Since the ODP is a relative measure, it is fairly robust, not overly sensitive to changes in the input data or to the details of the model calculations. That is, there are many uncertainties in calculating the numerator or the denominator of the expression, but most of these cancel out when the ratio is calculated. The nature of the halogen (bromine-containing halocarbons usually have much higher ODPs than chlorocarbons, because atom for atom Br is a more effective ozone-destruction catalyst than Cl.) The number of chlorine or bromine atoms in a molecule.
Molecular Mass (since ODP is defined by comparing equal masses rather than equal numbers of moles.) Atmospheric lifetime (CH3CCl3 has a lower ODP than CFC-11, because much of the CH3CCl3 is destroyed in the troposphere.) The ODP as defined above is a steady-state or long-term property. As such it can be misleading when one considers the possible effects of CFC replacements. Many of the proposed replacements have short atmospheric lifetimes, which in general is good; however, if a compound has a short stratospheric lifetime, it will release its chlorine or bromine atoms more quickly than a compound with a longer stratospheric lifetime. Thus the short term effect of such a compound on the ozone layer is larger than would be predicted from the ODP alone (and the long-term effect correspondingly smaller.)(The ideal combination would be a short tropospheric lifetime, since those molecules which are destroyed in the troposphere don’t get a chance to destroy any stratospheric ozone, combined with a long stratospheric lifetime.) To get around this, the concept of a Time-Dependent Ozone Depletion Potential has been introduced [Solomon and Albritton] [WMO 1991]: Loss of ozone due to X over time period T ODP(x,T) == ———————————————- Loss of ozone due to CFC-11 over time period T As T-*infinity, this converges to the steady-state ODP defined previously. The following table lists time-dependent and steady-state ODP’s for a few halocarbons [Solomon and Albritton] [WMO 1991] Compound Formula Ozone Depletion Potential 10 yr 30 yr 100 yr Steady State CFC-113 CF2ClCFCl2 0.56 0.62 0.78 1.10 carbon tetrachloride CCl4 1.25 1.22 1.14 1.08 methyl chloroform CH3CCl3 0.75 0.32 0.15 0.12 HCFC-22 CHF2Cl 0.17 0.12 0.07 0.05 Halon – 1301 CF3Br 10.4 10.7 11.5 12.5 —————————– Subject: 2.9) What about HCFC’s and HFC’s? Do they destroy ozone? HCFC’s (hydrochlorofluorocarbons) differ from CFC’s in that only some, rather than all, of the hydrogen in the parent hydrocarbon has been replaced by chlorine or fluorine. The most familiar example is CHClF2, known as HCFC-22, used as a refrigerant and in many home air conditioners (auto air conditioners use CFC-12).
The hydrogen atom makes the molecule susceptible to attack by the hydroxyl (OH) radical, so a large fraction of the HCFC’s are destroyed before they reach the stratosphere. Molecule for molecule, then, HCFC’s destroy much less ozone than CFC’s, and they were suggested as CFC substitutes as long ago as 1976. Most HCFC’s have ozone depletion potentials around 0.01-0.1, so that during its lifetime a typical HCFC will have destroyed 1-10% as much ozone as the same amount of CFC-12. Since the HCFC’s are more reactive in the troposphere, fewer of them reach the stratosphere. However, they are also more reactive in the stratosphere, so they release chlorine more quickly. The short-term effects are therefore larger than one would predict from the steady-state ozone depletion potential.
When evaluating substitutes for CFC’s, the time-dependent ozone depletion potential, discussed in the preceding section, is more useful than the steady-state ODP. [Solomon and Albritton] HFC’s, hydrofluorocarbons, contain no chlorine at all, and hence have an ozone depletion potential of zero. (In 1993 there were tentative reports that the fluorocarbon radicals produced by photolysis of HFC’s could catalyze ozone loss, but this has now been shown to be negligible [Ravishankara et al. 1994]) A familiar example is CF3CH2F, known as HFC-134a, which is being used in some automobile air conditioners and refrigerators. HFC-134a is more expensive and more difficult to work with than CFC’s, and while it has no effect on stratospheric ozone it is a greenhouse gas (though somewhat less potent than the CFC’s).
Some engineers have argued that non-CFC fluids, such as propane-isobutane mixtures, are better substitutes for CFC-12 in auto air conditioners than HFC-134a. —————————– Subject: 2.10) *IS* the ozone layer getting thinner? There is no question that the ozone layer over antarctica has thinned dramatically over the past 15 years (see part III). However, most of us are more interested in whether this is also taking place at middle latitudes. The answer seems to be yes, although so far the effect are small. After carefully accounting for all of the known natural variations, a net decrease of about 3% per decade for the period 1978-1991 was found. This is a global average over latitudes from 66 degrees S to 66 degrees N (i.e.
the arctic and antarctic are excluded in calculating the average). The depletion increases with latitude, and is somewhat larger in the Southern Hemisphere. Over the US, Europe and Australia 4% per decade is typical; on the other hand there was no significant ozone loss in the tropics during this period. (See, however, [Hofmann et al. 1996] for more recent trends which appear to show a decline in some tropical stations.) The depletion is larger in the winter months, smaller in the summer.
[Stolarski et al.] [WMO 1994] The following table, extracted from a much more detailed one in [Herman et al. 1993], illustrates the seasonal and regional trends in percent per decade for the period 1979-1990: Latitude Jan Apr Jul Oct Example 65 N -3.0 -6.6 -3.8 -5.6 Iceland 55 N -4.6 -6.7 -3.1 -4.4 Moscow, Russia 45 N -7.0 -6.8 -2.4 -3.1 Minneapolis, USA 35 N -7.3 -4.7 -1.9 -1.6 Tokyo 25 N -4.2 -2.9 -1.0 -0.8 Miami, FL, USA 5 N -0.1 +1.0 -0.1 +1.3 Somalia 5 S +0.2 +1.0 -0.2 +1.3 New Guinea 25 S -2.1 -1.6 -1.6 -1.1 Pretoria, S. Africa 35 S -3.6 -3.2 -4.5 -2.6 Buenos Aires 45 S -4.8 -4.2 -7.7 -4.4 New Zealand 55 S -6.1 -5.6 -9.8 -9.7 Tierra del Fuego 65 S -6.0 -8.6 -13.1 -19.5 Palmer Peninsula (These are longitudinally averaged satellite data, not individual measurements at the places listed in the right-hand column. There are longitudinal trends as well. A recent reanalysis of the TOMS data yields trends that differ in detail from the above, being somewhat smaller at the highest latitudes.
[McPeters et al. 1996]. ) It should be noted that one high-latitude ground station (Tromso in Norway) has found no long-term change in total ozone change between 1939 and 1989. [Larsen and Henriksen][Henriksen et al. 1992] The reason for the discrepancy is not known. [WMO 1994] Between 1991 and 1993 these trends accelerated. Satellite and ground-based measurements showed a remarkable decline for 1992 and early 1993, a full 4% below the average value for the preceding twelve years and 2-3% below the lowest values observed in the earlier period.
In Canada the spring ozone levels were 11-17% below normal [Kerr et al.]. By February 1994 ozone over the United States had recovered to levels similar to 1991, [Hofmann et al. 1994b] and in the spring of 1995 they were down again, to levels lower than any previous year other than 1993. [Bojkov et al. 1995] Sulfate aerosols from the July 1991 eruption of Mt. Pinatubo are the most likely cause of the exceptionally low ozone in 1993; these aerosols can convert inactive reservoir chlorine into active ozone-destroying forms, and can also interfere with the production and transport of ozone by changing the solar radiation balance in the stratosphere.
[Brasseur and Granier] [Hofmann and Solomon] [Hofmann et al. 1994a] [McCormick et al. 1995] Another cause may be the unusually strong arctic polar vortex in 1992-93, which made the arctic stratosphere more like the antarctic than is usually the case. [Gleason et al.] [Waters et al.] In any event, the rapid ozone loss in 1992 and 1993 was a transient phenomenon, superimposed upon the slower downward trend identified before 1991. —————————– Subject: 2.11) Is the middle-latitude ozone loss due to CFC emissions? That’s the majority opinion, although it’s not a universal opinion.
The present trends are too small and the atmospheric chemistry and dynamics too complicated to allow a watertight case to be made (as has been made for the far larger, but localized, depletion in the Antarctic Ozone hole; see Part III.). Other possible causes are being investigated. To quote from the 1991 Scientific Assessment published by the World Meteorological Organization, p. 4.1 [WMO 1991]: The primary cause of the Antarctic ozo …