In this section, we discuss what are the different types of short-term, seasonal, interannual, and long-term variabilities controlling ozone amounts in the upper stratosphere. We will see that variability can be driven by changes in the photochemistry or by changes in dynamical or transport processes. The response to changes in the upper stratosphere is, however, governed by photochemical processes. In the next section (5), we will examine the lower stratosphere where dynamical (transport) processes greatly influence the ozone distribution and its variability.

4.1 Short-Term Variability in the Upper Stratosphere

There are several types of short-term variability that affect ozone photochemical process rates in the upper stratosphere. These include diurnal variations, variations in solar ultraviolet radiation, temperature driven variations, and particle precipitation events that originate from electromagnetic storms on the Sun.

4.1.1 Diurnal variations -- In the upper stratosphere, above 40 km, where PRTs are less than 1 day, variations in ozone occur with the daily rising and setting of the Sun. These variations are usually termed "diurnal" because they happen each day.

(a) Cycle of ozone creation and destruction in the upper stratosphere -- To illustrate this, we consider the steps involved in a simple, pure oxygen description of upper stratospheric ozone chemistry.

Step 1: Photodissociation of oxgyen molecules into free oxygen -- Ozone is created when ultraviolet light of wavelength less than 240 nm strikes an atmospheric oxygen molecule splitting it into two oxygen atoms

O2 + hf ( l < 240nm) --> O + O.

where h is Planck's constant, f is the frequency of radiation, and l is wavelength of the radiation. Wavelength and frequency are related through the c, speed of light, as c=f l, where c=3.00x 108 meters per second.

This process is known as photodissociation or photolysis. The resultant O (oxygen) atoms undergo numerous collisions with N2 (nitrogen) and O2 (oxygen) molecules. Collision with an O2 molecule can lead to a combination of O2 and O which has some excess energy of collision. The lifetime of this complex is short and most of the time it will simply fly apart without any additional reaction taking place.

Step 2: Creation of stable ozone molecule in a three body reaction -- Occasionally, however, another molecule (denoted M) will collide with the complex and absorb the excess energy, leaving intact a stable O3 (ozone) molecule. This is a three body reaction which converts atomic oxygen to ozone

O + O2 + M --> O3 + M

This three body reaction is the usual way that an ozone molecule forms. Why should this be so? From basic physics, we know that any collision must conserve momentum and energy simultaneously. The easiest way to view the collision of two molecules, such as O and O2, is to use the center-of-mass coordinate system. The approaching molecules are viewed in a coordinate system that moves with the center of mass of the two particles. In this coordinate system, the two particles collide and coalesce into one particle which is not moving. Momentum is conserved. The kinetic energy in this coordinate system must be zero after the collision because the single remaining particle is at rest at the center of mass. The kinetic energy has been converted into internal energy of the combined particle O·O2. If this amount of energy exactly coincides with an internal energy level of O3, the resulting O3 molecule is stable. This is highly unlikely. More likely is that a third body collides with the excited O·O2 complex whose energy does not exactly coincide with the internal energy level of O3. The excess energy can then be distributed between the relative kinetic energies of the two remaining particles, the O3, and the third body. In this three-body reaction, it is the third molecule that greatly increases the chances of forming a stable ozone molecule by redistributing relative kinetic energies in such a way that a stable ozone molecule is formed. This is why most ozone is formed in three body reactions.

The time constant in the upper stratosphere for converting O atoms and O2 molecules to ozone ranges from a few seconds at 40 km up to about 100 seconds at 50 km. The time constant increases with increasing altitude as the densities of both O2 and M decrease.

Step 3: Photodissociation of ozone molecule by UV light into O2 and O -- Once formed, the O3 molecule is stable until it is split apart once sunlight is present by solar ultraviolet radiation of wavelength shorter than about 310 nanometers (in a process called photodissociation) back to O2 and O by light. The reaction is given by

O3 + h f (l < 310nm) --> O2 + O

The PRT for this process is about 1000 seconds in the upper stratosphere in the presence of sunlight.

Step 4: Recombination reaction of O and O3 back to O2 -- Finally, the life cycle of ozone is completed by the recombination reaction of atomic oxygen with an ozone molecule to reform the original molecular oxygen in a reaction given by

O + O3 --> O2 + O2

The time constant is about one day at 40 km decreasing to a fraction of a day at 50 km. This reaction can occur exactly as given or it can be catalytically speeded up by processes involving the oxides of hydrogen, nitrogen, chlorine, and bromine, which we do not consider here (see Chapter 5 for details).

(b) Summary of cycle -- The picture of the upper stratospheric ozone chemistry we have described can be summarized with the following concepts. First, UV radiation of wavelength less than 240 nm splits apart molecular (O2) molecules, creating atomic oxgyen. This atomic oxygen rapidly reacts with another O2 molecule to form ozone. After sunset, this "photolysis" of O2 is shut off, and all the available O atoms are converted to O3 (ozone) molecules in three body reactions that occur within a few seconds to minutes. The (small) amount of atomic oxygen falls off quickly while the amount of ozone rises slightly. As the Sun rises, some of the ozone molecules are photodissociated by UV radiation with wavelengths less than 310 nm. Finally, O3 molecules can react with O atoms to reform O2 molecules. Both the photolysis of O2 and the reaction of O with O3 are relatively slow processes compared to O3 photolysis and the reaction of O with O2 molecules.

While ozone is produced and destroyed on a time scale of one day or less at altitudes greater than 40 km altitude, this does not in itself lead to a diurnal variation of ozone. This is because when sunlight is present, production and loss are balanced (as given by the steady state approximation in the upper stratosphere). When the sun sets, both production and loss are turned off and the ozone concentration increase slightly as O atoms are converted to ozone molecules in three body reactions involving O, O2, and M. It is the conversion of atomic oxygen to ozone after sunset followed by its reconversion to atomic oxygen at dawn that is primarily responsible for diurnal variations of ozone in the upper stratosphere.The diurnal variation of ozone in the upper stratosphere results from the conversion of O to O3 after sunset by three-body reactions with O2 and M and the reconversion of O3 to O by sunlight at dawn.

The ratio of the concentration of O to that of O3 is governed by the rapid photochemical reactions which interconvert O and O3. At steady state, the rate of creation of ozone from three body collisions equals the rate of photodissociation of ozone by sunlight. Written out as an equation, we have

k(O,O2)·[O]·[O2]·[M] = J(O3)·[O3]

where the brackets [] around a chemical symbol indicate the concentration of the chemical in molecules/cm3; k(O,O2) is the reaction rate coefficient for the reaction of O with O2 in cm6molecule-2sec-1; and J(O3) is the photodissociation coefficient for O3 in molecule-1sec-1. (See Chapter 5 for a description of reaction rate coefficients and photodissociation coefficients.)

This equation can be solved for the ratio of [O] to [O3]

[O]/[O3] = J(O3)/(k(O,O2)·[O2]·[M])

This ratio is extremely small in the lower stratosphere but increases with increasing altitude as both the [O2] and [M] in the denominator fall off exponentially. By 50 km, the daytime ratio of [O]/[O3] is a little less than 0.1. Thus, ozone at 50 km will increase by about 10% at night and then decrease again as sunlight returns.

4.1.2 Solar UV variations -- We have devoted considerable space above to the theory behind the relatively minor diurnal variation. We can, however, use the same theory we have developed to examine other possible variations in upper stratospheric ozone. One of these is the variation of ozone in response to variations in the ultraviolet output of the Sun.

The output of solar ultraviolet radiation is influenced by magnetically active regions on the Sun. These occur sporadically in sunspots, which are cooler regions on the Sun (hence their "dark" appearance compared to the rest of the Sun). The magnetic field is especially strong in sunspot regions. Sunspots are easily observable with a simple telescope (though one must never look at the Sun through a telescope without filters specifically designed for viewing the Sun). They occur in quasi-regular 11-year cycles (see section on interannual variability for further discussion). The period when sunspot activity is at its greatest during the 11-year cycle is a solar maximum, while the period when sunspot activity is at its least is a solar minimum. During solar maxima, sunspots tend to occur in clusters on one side of the Sun. The Sun rotates with a 27-day period, so the ultraviolet radiation from the Sun is modulated by this rotational period. Occasionally, two active sunspot regions will occur on opposite sides of the Sun and a 13-day modulation of ultraviolet radiation occurs.

These modulations in the solar ultraviolet output have a direct effect on ozone photochemistry and even ozone column amount (see Chapter 9). The production rate of ozone can be affected by changes in J(O2), the photodissociation rate of O2. The loss rate can also be affected by changes in J(O3), the photodissociation rate of O3. Such changes produce atomic oxygen, leading to loss through the O + O3 reaction and the O + oxides of hydrogen, nitrogen, chlorine, and bromine reactions (see Chapter 5). The effect on ozone in the photochemical region of the upper stratosphere then depends on the relative importance of variations in J(O2) and J(O3).

The effect of sunspots on solar ultraviolet radiation output depends strongly on the particular wavelength of ultraviolet radiation. Table 8.01 gives the change in the Sun's ultraviolet energy output from solar maximum to solar minimum in a typical sunspot cycle as a function of wavelength (from Dessler et al., 1998). The UV output of the Sun increases as sunspot activity increases.

Table 8.01
Approximate variability of solar ultraviolet radiation over a solar cycle as a function of wavelength
200 nm
220 nm
270 nm
300 nm

We see in Table 8.01 that the wavelengths shorter than 240 nm that contribute to the breakdown (photolysis) of O2 molecules show variations of 4-8% from maximum to minimum. The photolysis of O3 in the upper stratosphere is dominated by UV flux with wavelengths from 240 nm to 320 nm. The output (flux) of these particular wavelengths of UV radiation vary 1% or less from solar maximum to minimum. Thus, the production rate of ozone (via O2 photolysis and denoted by J(O2) in our notation) is more sensitive to variations in solar UV output than is the loss rate of ozone (via O3 photolysis and denoted by J(O3) in our notation). The amount of UV radiation increases with increasing sunspot activity, reaching a maximum at solar maximum. The increases occurs principally in the shorter wavelength UV radiation (less than 240 nm), so that there is more photolysis of O2 and hence more production of O3 than there is photolysis (loss) of O3. This means that the production term of ozone increases and overall O3 concentrations increase when solar activity is at maximum. The result is a direct, inphase variation of upper stratospheric ozone concentrations with the UV flux changes from sunspot activity.

Figure 8.03 shows a plot of ozone mixing ratio at 1 mbar (about 48 km altitude) versus time for a 4 month period in 1980. (For explanation of mixing ratio, see Chapter 3.) This period was near a solar maximum in sunspot activity. Sunspots were characteristically clustered on one side of the sun. The plot of ozone mixing ratio over the four month period shows a succession of four peaks separated by 27 days each, corresponding to the solar rotation period. The amplitude of these oscillations is about 1.5% of the ozone amount. Increases in ozone mixing ratio occur when the sunspot active regions of the Sun are facing Earth. Again, this occurs because sunspot activity corresponds to increased UV flux in those wavelengths where increased photolysis of O2 and hence increased production of O3 occurs.

4.1.3 Temperature driven fluctuations -- Although the upper stratosphere is dominated by photochemical production and loss processes, dynamical variations also produce short-term variations. Planetary waves (see Chapter 6) propagating upward from the lower atmosphere grow in magnitude as altitude increases. These waves lead to variations in temperature which then affect the photochemical reaction rates, because the reaction rate coefficients that determine photochemical loss rates are temperature dependent.

The mechanism for the temperature dependence is described in some detail in Chapter 5. The key is the temperature dependence of the reactions involved in ozone loss. The rate coefficient for the reaction O + O3 --> O2 + O2 as a function of temperature has been determined from laboratory measurements. It can be approximated by

k(O,O3) = 8.0 x 10-12 e-2060/T cm3molecule-1sec-1

where T is the temperature on the Kelvin scale.

This temperature dependence means that at T=250K(-23C), a change of 10 degrees in the temperature results in a greater than 30% change in the reaction rate coefficient.

We are only been considering pure oxgyen photochemistry, though we've already alluded to the fact that ozone photochemistry can be speeded up by catalytic reactions involving various oxides of nitrogen, hydrogen, chlorine, and bromine. The nitrogen oxide catalytic cycle has a similar strong dependence on temperature as the pure oxygen case. The hydrogen, chlorine, and bromine oxide catalytic cycles each have a much weaker temperature dependence. These catalytic cycles are discussed in detail in Chapter 5.

The temperature dependence described above is such that the loss rate increases with increasing temperature. This means ozone amounts will decrease when temperatures increase and are thus anticorrelated with temperature. This effect will be strongest in winter when wave activity propagating into the stratosphere is strongest. An example is shown in Figure 8.04 for the northern hemisphere winter period of 1984-85.

4.1.4 Short-term events (solar storms) -- The principal type of solar events in which an effect on ozone has been documented is the solar proton event. Solar storms lead to ejection of large amounts of high energy protons that can penetrate the Earth's magnetic field near the poles. These protons penetrate into the atmosphere, typically to the 40 to 80 km layer, causing ionization of air molecules. As the ionized particles recombine, they produce nitrogen and hydrogen oxides which can affect ozone through the NOX and HOX catalytic cycles discussed in Chapter 5. The effects are short lived because the hydrogen oxides which cause the primary ozone loss recombine within hours. The effects of nitrogen oxides can persist for several months.

4.2 Seasonal Variability in the Upper Stratosphere

As we saw above in section 4.1.3, there is an anticorrelation between temperature and ozone concentration. In a similar manner, there are seasonal variations in upper stratospheric ozone concentrations that are driven by the seasonal change in temperature.

The driving force in creating seasonal temperature variations at the surface and throughout the atmosphere, including the upper stratosphere, is the seasonal change in the elevation angle of the sun. The mechanisms differ, though, depending on altitude and latitude.

Earth's surface absorbs the shortwave visible radiation from the sun that passes through the atmosphere mostly uninterrupted (see Chapter 4). This energy is then reemitted as thermal longwave (infrared) radiation. The lower atmosphere is warmed in turn by the absorption and reemission of this infrared (IR) radiation. The progressive upward reemission of IR radiation warms the atmosphere, but as air density drops off, the efficiency of heat transfer by reemission decreases, so less warming occurs and temperature falls off with altitude. This fall off in temperature is called the lapse rate (see Chapter 2) and depending on the particular value of it, the atmosphere is said to be stable, neutral, or unstable. Convective instability occurs when the lapse rate is especially steep (so that the potential temperature falls off with height, see Chapter 2 for explanation) that turbulent mixing results as large air parcels overturn and mix.

In the upper stratosphere, absorption of solar UV radiation by ozone itself provides the source of heating to the atmosphere. In summer, when the sun is nearly overhead, this heating is maximized, leading to a temperature maximum. The opposite occurs in winter when the sun is much lower in the sky. The maximum temperature occurs near the summer solstice (June 21 or day number 172) and the minimum temperature occurs near the winter solstice (December 21). Ozone amount is inversely related to the temperature: it reaches a minimum when the temperature is at a maximum, and it reaches a maximum when the temperature is at a minimum. Figure 8.05 demonstrates this inverse relationship with a plot of seasonal variation of ozone and temperature at 1 mbar (about 48 km altitude) for 1988 for the 40-45°N latitude belt. This inverse relationship is due simply to the photochemical loss reactions, which are temperature dependent decreasing as temperature increases.

4.2.1 Radiative equilibrium -- Another aspect of upper stratospheric warming due to ozone implied by Figure 8.05 is the radiative equilibrium temperature. Radiative equilibrium is established when the temperature of an air parcel is determined by a balance between heating from absorption of solar energy and cooling to space. It is also determined by any heating from adjacent warmer air parcels and cooling to adjacent cooler air parcels. In the case of the upper stratosphere, the heating is primarily from absorption of UV radiation by ozone. The cooling occurs through infrared emission by carbon dioxide with a small contribution from water vapor. The cooling by carbon dioxide at these altitudes is through infrared radiation reimission to space. Interestingly, this means that as carbon dioxide increases and the surface is expected to warm through the atmospheric greenhouse effect, the stratosphere is actually expected to cool.

The presence of ozone results in a maximum temperature that is 5-10K cooler than the corresponding radiative equilibrium temperature, while the minimum temperature is 5-10K warmer.

4.2.1 Midlatitude ozone variability -- In addition to displaying altitude variations, ozone varies by latitude. The variations in latitude arise from the fact that ozone is created in the tropics, whence it undergoes an equator to pole circulation. The circulation is driven by temperature differences. In the hemisphere where it is winter, bigger temperature differences lead to a stronger ozone circulation. This seasonal equator-to-pole circulation is known as the Brewer-Dobson circulation (see Chapter 6).

In the midlatitudes ozone in the upper stratosphere has a regular seasonal variation with a minimum in the early summer and a maximum in the early winter. The opposite of that for temperature. This is associated with the same temperature dependent chemistry in Section 4.1.3. Figure 8.05 demonstrates this inverse relationship. The magnitude of the seasonal variation of ozone and temperature in the upper stratosphere is made somewhat smaller than it would otherwise be because of a negative feedback effect. In summer when the temperature is a maximum, less ozone is available to absorb solar radiation.

4.2.2 Equatorial ozone variability -- Equatorial latitude seasonal variations of ozone in the upper stratosphere are more complicated. The Sun passes overhead twice during the year at the equator. This would lead us to expect a semiannual oscillation (two peaks during the year) in the temperature and hence in ozone amounts. This is indeed the case as seen in Figure 8.06. However, the semiannual oscillation of tropical temperatures is not symmetrical and the timing is not quite in phase with the passage of the Sun. This is an indication of the importance of stratospheric dynamics in modifying the temperatures which would occur if the atmosphere were in simple radiative equilibrium.

4.3 Interannual Variability in the Upper Stratosphere

Now that we have a framework for understanding short-term changes in upper stratospheric ozone, we can go on to examine some of the causes for year-to-year variations. These are variations that we previously referred to as interannual variability.

4.3.1 Quasi-Biennial Oscillation (QBO) -- The QBO is of only minor importance to the variability of the upper stratosphere. Its effects on ozone amount and distribution are primarily in the lower stratosphere. This is explained in more detail in Section 5.3.1. The relationship of QBO and stratospheric ozone is also discussed in Chapter 9.

4.3.2 The 11-year solar cycle -- We have already discussed the potential impact of solar activity on short-term variations of ozone. These occurred through the 27-day solar rotation of active sunspot regions. The frequency sunspot occurrence goes through an 11-year cycle. During solar maximum, numerous active regions occur and certain wavelengths of ultraviolet radiation are enhanced. During solar minimum, the active regions disappear, solar ultraviolet radiation is at a minimum, and the 27 day cycle disappears. Ozone concentrations in the upper stratosphere increase with increased UV radiation and decrease with decreased UV radiation, as discussed in section 4.1.2.

Because of the 11-year variations in the output of solar ultraviolet radiation we expect an 11-year variation in upper stratospheric ozone through the same mechanism responsible for the 27-day variation in ozone. This is indeed what is observed. Increased solar ultraviolet radiation during the maximum phase of the solar cycle (see Table 8.01) leads to more production of ozone via the photolysis of O2. In the photochemical steady state of the upper stratosphere, this leads to an increase in the concentration of ozone. Figure 8.07 shows a plot of upper stratospheric ozone at 2 mbar for a 12-year period, 1979-1991, that includes a full solar cycle. The data is based on measurements from the SBUV instruments aboard the Nimbus-7 and NOAA-11 satellites.

The data shown in Figure 8.07 has had the mean value removed as well as the mean trend and seasonal variation. The resulting deviations show a clear signature of the 11-year solar cycle. Two solar cycle maxima, one around 1979-80 and another in 1989-90, correspond to times of largest positive deviations in ozone mixing ratios, while the one solar minimum in the period around 1984-85 corresponds to the time of largest negative deviation in ozone mixing ratio.

We must express some caution because the data captures only one solar cycle (maximum-to-maximum). However, the data are entirely consistent with what would be expected from the variation in solar ultraviolet radiation output.

4.3.3 Volcanic eruptions -- Explosive volcanic eruptions can inject large amounts of sulfate aerosols well into the stratosphere. The largest eruptions, such as the June 1991 eruption of Mount Pinatubo in the Philippines, can inject material into the atmosphere as high as 30-35 km. This is almost to the level of the upper stratosphere. There may be aerosol effects on upper stratospheric ozone, but such events are rare. Most volcanic eruptions inject material no higher than the lower stratosphere, and it is here that effects on ozone are centered. The effects on lower stratospheric ozone are often significant. These are discussed in Section 5 of this chapter.

Volcanic eruptions could also affect the ozone amounts in the upper stratosphere indirectly by perturbing the dynamics of the stratosphere which, in turn, affect the temperature of the upper stratosphere. This occurs through temperature dependent chemistry. This effect is difficult to evaluate, and is not considered to be important in the upper stratosphere.

4.4 Long-Term Variability in the Upper Stratosphere

Other influences on ozone concentrations in the upper stratosphere work on longer time scales than those considered so far. These are referred to as long-term variability.

4.4.1 CFC impact on upper stratospheric ozone -- The most important influence on the long-term variability of stratospheric ozone is the introduction of significant amounts of chlorine into the upper stratosphere from industrially produced chlorofluorocarbon (CFC) compounds (see Chapter 10 for more detailed information). The initial predictions of effects of chlorine from CFCs indicated that the change in ozone would be a maximum in the upper stratosphere near 40 km altitude. This change was predicted to result from a relatively simple set of reactions that form the chlorine, bromine, nitrogen, and hydrogen catalytic cycles for ozone destruction. Observations over the past 20 years indicate that these decreases are indeed occurring (see Chapter 9).

4.4.2 Unresolved longer cycles -- When analyzing ozone data for the long-term trend, we must always be aware of the shortness of our "long-term" data set. It is possible that our data set is simply not long enough to resolve some natural cyclic pattern. There could even be a noncyclic factor unresolved by our limited data set. Some unknown factor contributing to a variation with a period longer than the period of our record might cause us to conclude erroneously the existence of a long-term trend in the data.

Solar activity can be used as an example of unresolved longer cycles. We have discussed the short term 27-day solar rotation and the 11-year sunspot cycle. The magnitude of the sunspot number during solar maximum is also variable. Among other variations it has been observed to exhibit a 90-year cycle called the Gleissberg cycle, as shown in Figure 8.08. There are three of these cycles with peaks in about 1780, 1870, and 1960. Previous to this, during the 1600's, virtually no sunspots were observed in a period that is referred to as the Maunder minimum.

Such long-term periodicity in an external forcing produces long-term cyclic variations in ozone. A data set of only 20 years would not capture the effects due to the Gleissberg cycle. Such possibilities must always be kept in mind when trying to identify long-term trends.

We have no way to be certain that there are no major unresolved long-term variations. The best we can do is to try to understand all of the variations observed and attribute them to physical processes.

4.5 Relationship to Total (Column) Ozone Variability

Our discussion so far has been about variations in ozone in the upper stratosphere. On average, the amount of ozone above 40 km altitude is about 3% of the total ozone, and the amount above 30 km altitude is about 20% of the total ozone. Thus, those variations occurring above 40 km, such as the diurnal variation, have little measurable effect on total ozone column (10% of 3% is about .3% or three parts in a thousand variation). Variations occurring over the broader range of altitudes above 30 km can have a little larger effect. However, to produce major variations in total ozone, it is really necessary to alter the ozone in the lower stratosphere where the density reaches a maximum. We turn to ozone variability in the lower stratosphere in the next section.