We have seen that variations in the ozone concentration in the upper stratosphere occur on a variety of time scales. Because the ozone in the upper stratosphere is only a small fraction of the total ozone, we have argued that the variations in the total column amount of ozone must be dominated by changes in the lower stratosphere. We will now consider variations in the lower stratosphere in more detail starting with the shortest time scales.
The time scale for ozone loss becomes very long in the lower stratosphere (as much as 1 year). This is much longer than the time to move air from one place to another. The time scales increase because the ozone in the lower stratosphere is shielded from solar photodissociation by the ozone in the middle and upper stratosphere. This means that in the lower stratosphere, the amount of ozone at any one place depends on where the air came from and how much ozone was contained in that air. Variability at these altitudes is more dependent on transport processes than on photochemical ones.
5.1.1 Passage of weather systems -- The motion of ozone is directly associated with the dynamics of motion in the atmosphere. The usual wintertime flow in the stratosphere of the midlatitudes is from west-to east (westerly). This reverses in the summertime, becoming east-to-west (easterly). The passage of a wintertime weather systems is equivalent to the superposition of a secondary air flow on top of the overall westerly flow (see Chapter 6).
In addition to the overall flow of air, the motion of individual air masses can be followed if we know the particular temperature, moisture content, and amounts of any trace gases (e.g., ozone) in different air masses. The motion of such characteristics is called advection. Thus, we refer to warm air advection or cold air advection or moisture advection. Like these other types of advection, ozone advection shows the effect of weather systems. This overall way of moving ozone around is referred to as a transport process. It differs from photochemical processes that actually create and destroy ozone. Transport merely redistributes ozone from place to place.
When a high pressure system is generated in the troposphere, such as you would see on a weather map, the motion of air within the high is downward. Air in the tropospheric column sinks. Such sinking motion results in compression and heating of the air. The motion of stratospheric air above this tropospheric weather feature is in the opposite sense. In the case of the tropospheric high pressure system, the motion of the stratospheric air above the high is upward. Air in the stratospheric column rises.
In our example, above the tropospheric high pressure system, rising air is brought into the stratospheric column near the bottom of the stratosphere (at the tropopause) and is removed from the column at higher altitudes. Since ozone mixing ratio in the lower stratosphere increases with increasing altitude, the air brought into this column has less ozone than the air being removed and ozone in the column decreases. Once this high pressure system is formed and moves, it carries the low ozone column along with it. This is why high pressure systems are associated with decreased ozone amounts. The effects of a low pressure system are the opposite, and ozone amounts are increased. The result is that a map of total ozone amounts strongly resembles a weather map (see Chapters 2 and 6).
These variations are mostly just temporary redistributions of ozone with no net gain or loss. The photochemical time constants are slow in the lower stratosphere compared to these transport processes. Thus, transport processes are responsible for most of the observed variations in lower stratospheric ozone. A decrease in a high pressure region is balanced by an increase in surrounding regions. We address the implications for seasonal variability in the next section.
In the upper stratosphere, the seasonal variability of ozone concentrations are driven by the seasonal variation in temperature acting through temperature dependent photochemistry. In the lower stratosphere, photochemistry is much slower, owing to reduced UV flux, and transport of ozone by large-scale atmospheric motions (known as circulations) becomes important.
We saw in Chapter 2 that midlatitude total column ozone exhibits a strong seasonal variation. This variation has a maximum in late winter/early spring and a minimum in late fall. This variation is different from what is observed in the upper stratosphere and from what we would expect from pure photochemical theory.
The primary driver of the springtime maximum in total ozone is the transport of ozone rich air towards the poles during the winter. Why is ozone transported to the poles during winter? The primary circulation of the lower stratosphere is upward in the tropics and downward/poleward at mid- and high latitudes. This circulation is driven by the dissipation of wave disturbances propagated upward from the troposphere. (See Chapter 6 for details.)
Large-scale wave disturbances are generated in the troposphere by flow over mountain ranges and across the temperature contrast of land-sea interfaces. These waves have a spectrum of spatial and temporal scales. The stratosphere acts like a filter for these waves. Only the larger scale waves can propagate into the stratosphere and those do so only in winter. The stratospheric mean flow is westerly in winter and easterly during summer. The large-scale waves (Rossby waves) propagate into the winter westerlies but not the summer easterlies.
The dissipation of these large-scale waves in the stratosphere leads to a wintertime downward/poleward circulation. This circulation carries with it ozone which was produced in the tropical photochemical source region. This ozone rich air is moved to high latitudes where, during winter, the photochemical time constant for ozone loss is very long (about one year). The time constant is long because the ozone is effectively shielded from ultraviolet radiation by the ozone above, and because the sun angle is quite low in the high latitude winter. The process occurring is much like storing the ozone in the lower stratosphere where it is safe from destruction by photochemical processes. This accumulation of ozone rich air over the entire winter period leads to a maximum in the springtime.
As summer approaches, large-scale waves stop propagating into the stratosphere. The Sun's position in the sky climbs higher. Photochemistry can now occur, causing the ozone concentration to decrease until a minimum is reached in late fall. The accumulation process then begins again as winter returns to the stratosphere.
As in the upper stratosphere, there are factors that produce year-to-year variations in lower stratospheric ozone. These are the interannual variations, chief among which are the QBO and the ENSO, both of which are described in more detail below. There are also other variations with irregular periods, such as volcanic eruptions, leftover effects from atmospheric nuclear testing in the 1950s and 1960s, and chaotic variations due to internal variability in stratospheric dynamics.
5.3.1 Quasi-Biennial Oscillation (QBO) -- The Quasi-Biennial Oscillation is an oscillation in the average zonal winds in the tropical stratosphere. Roughly every 27-30 months, the tropical stratospheric winds in the 10 mb to 100 mb altitude range are observed to shift from westerly to easterly and then back again. Weather systems in the tropical lower stratosphere reflect this shift.
The QBO develops as a result of disturbances, or waves, in the tropical troposphere propagating vertically into the lower stratosphere. These disturbances interact with the average wind in such a way as to cause the average wind to regularly reverse direction. We can see this shift in Figure 2.16 of Chapter 2. When we look at the winds versus height, we see that areas of easterly and westerly winds are stacked on top of each other. As we look forward in time, we can see that these winds descend at a rate of ~1 km/month. If we were to observe the winds from a fixed point in the tropical stratosphere, we would see that over the course of about 2.5 years, the winds will have gone through a complete cycle. For instance, if the winds were from the east when we started our observation, over time they would they weaken, then reverse to westerly winds. The westerly winds would gain strength for a period and then weaken again, finally reversing to easterly winds once again.
The QBO is fascinating in itself, but what does it have to do with ozone? In the last section, we saw that transport is a major component of ozone change in the lower stratosphere. The QBO not only changes the circulation in the tropics, but also indirectly causes changes at middle and high latitudes as well. The waves that cause the reversal of the tropical winds do so by exerting a drag on the flow. During one phase of the QBO, this drag induces a circulation through the stratosphere. This circulation enters the lower stratosphere in the tropics and leaves the middle stratosphere in the extratropics. During the other phase of the QBO, the induced circulation reverses itself. Momentum is thus exchanged between the tropics and the extratropics. Air is also transported between the two latitude regions. These air masses have different ozone concentrations.
The induced meridional (in the latitude/altitude plane) circulation moves ozone poor air into the tropics and ozone rich air out of the tropics during one phase of the QBO and then reverses this process during the other. Ozone in the tropics thus decreases and then increases as we go through the QBO cycle. These induced circulations reach to midlatitudes, increasing or decreasing the ozone concentration there in the opposite phase to what is happening in the tropics. This variation of the phase of the ozone response to the QBO is seen in statistical model analyses of ozone variations with latitude as described in Chapter 9.
The simple explanation for QBO variations given above would imply that ozone is conserved during this cycle and the globally integrated amount would not vary. However, this is not the case. The globally averaged total amount of ozone also has a QBO signal (see Chapter 9) which is opposite to the phase at the equator. This is because the induced circulation reaches up into the zone of rapid photochemistry. When the circulation is upward at the equator, air is pulled out of the photochemical reaction zone and deposited in the midlatitudes. This photochemical zone acts like an infinite source because of continuous balancing of production and loss. It is limited only by its own concentration. This air pulled out of the photochemical zone leads to an increase in the overall amount of ozone.
5.3.2 The El Niño-Southern Oscillation (ENSO) -- Large variations in equatorial Pacific Ocean surface temperatures are known as the El Niño-Southern Oscillation (ENSO). Normally, the waters of the Eastern Pacific Ocean near South America are quite cool as a result of upwelling ocean currents. The tropospheric trade winds normally traverse from east to west in this region. However, in an El Niño period, the trade winds weaken, allowing the warmer waters of the Western Pacific to migrate eastward. This change in sea surface water temperature causes large-scale shifts in the global circulation patterns in the troposphere and lower stratosphere. This in turn affects the transport of ozone in these regions. This oscillation is very irregular, with a period of 4-7 years between episodes.
Ozone variations in response to ENSO should be strongest in the tropics with opposite signs at different longitudes. The zonal mean ENSO variation should be very small. ENSO weather patterns reach into the midlatitudes, as is well known from the multitude of news stories related to the recent major El Niño (1997-98). Ozone effects also reach into midlatitudes, although the effects are hard to separate from other variations.
5.3.3 Volcanic eruptions -- Explosive volcanic eruptions can inject large amounts of material directly into the lower stratosphere. There are a number of ways in which this material can have an effect on the ozone content of the lower stratosphere and on the total ozone content. These include
Determining what part of the ozone variation after a volcanic eruption is due to that eruption is a difficult problem. The chemical effect of an eruption is dependent on the chlorine content of the stratosphere because of the limited amount of chlorine available for conversion on sulfate surfaces. Thus, the Agung eruption in 1963 should have had less impact on ozone loss than that of El Chichon in 1982, since stratospheric chlorine content had increased between the eruptions.
The chemical effect is also strongly dependent on temperature. The largest effect is in the winter lower stratosphere where cold temperatures allow the sulfate particles to grow and increase their surface area. Any dynamical effects due to heating of the aerosols will be difficult to separate from other variations in the dynamics.
5.3.4 Nuclear testing -- During the late 1950s and early 1960s, extensive atmospheric testing of nuclear bombs was carried out. The temperature within the fireball of a nuclear explosion is sufficient to convert atmospheric N2 and O2 into nitrogen oxides. The fireball could reach altitudes in the middle stratosphere and deposit the nitrogen oxides, which then could participate in catalytic ozone destruction reactions.
The question of whether nuclear testing caused a measurable decrease in ozone has been argued for a long time. It is unfortunate that the amount of ozone data available in the late 1950s and early 1960s is very small. The earliest ozone trend paper (Komhyr et al., 1971) showed an increase in total ozone during the 1960s from the available station data. It has been suggested that this was a recovery from the nuclear testing at the beginning of that period, but there is no way of knowing for sure.
5.3.5 Chaotic variations -- One final source of interannual variability is the vagaries of the meteorology of the stratosphere. The atmosphere is a nonlinear system which is known to undergo chaotic variations. These variations lead to variations in ozone through the dynamical influence on wintertime ozone transport and through temperature dependent chemistry.
One of the primary reasons for the study of ozone variability that we have described thus far is the determination of long-term trends. We are particularly interested in any long-term trend which may be caused by chemical changes in the atmosphere due to human activities. Once we know cyclical variations in ozone, we can begin the process of removing these signals in order to determine any long-term ozone trend, including those trends caused by human activities. The theory behind the prediction of a long-term change in ozone due to human activities has been discussed in detail in several previous Chapters. In identifying long-term trends, however, as in the case of the upper stratosphere, it is important to know whether or not the data set is long enough to determine the difference between real trends and cyclic variations with periods longer than the data set.
5.4.1 CFC impact on lower stratospheric ozone -- Chlorofluorocarbons (or CFCs) were developed in 1928 as a benign and inert chemical compound for refrigeration, replacing toxic and flammable refrigerants like ammonia. CFCs quickly gained enormous usage over the next several decades as coolants and spray propellants. Seventy years later, manmade CFCs today represent the most important influence on the long term variability of lower stratospheric ozone. They also represent the most serious threat to the the ozone layer, which shields the Earth's surface from biologically destructive ultraviolet light emitted by the sun. Chapters 1 and 10 give additional background on CFCs, while Chapter 11 discusses in great depth the "ozone hole" problem that has developed over Antarctica as a result of CFCs.
How do CFC compounds impact lower stratospheric ozone? In brief, the stable CFC molecule rises into the lower stratosphere, where intense UV light breaks down the molecule, liberating chlorine. This chlorine is then free to react in catalytic cycles to destroy ozone, especially in the presence of polar stratospheric clouds (PSCs). These catalytic cycles represent more complicated photochemistry than the simple oxygen chemistry discussed so far. Chapters 5 and 11 detail these reactions.
Consequences on lower stratospheric ozone have been dramatic at certain times of the year over Antarctica. It is there that exceptionally cold winter conditions permit widespread PSC formation. Total column amounts of ozone in spring over the Antarctic have fallen by over 50 percent (see Chapter 1). Overall in the lower stratosphere, ozone concentrations have declined by smaller amounts over the last 20 years (see Chapter 9).
The primary driver for expected photochemical change is chlorine from CFC's. They were calculated to induce a nearly linear downward trend in the total global ozone amount starting in 1970 and continuing until about 1995. Because the provisions of the Montreal Protocol (see Chapter 10) have begun to take effect, this linear decrease in ozone is now predicted to flatten out and eventually turn around. Figure 8.09 shows a model calculation for the effect on ozone of increasing chlorine, the solar cycle, and volcanic activity. The model calculations project into the future using assumptions about the release of CFCs and other source gases such as methane and nitrous oxide.
In nearly all studies of long-term ozone trends it has been assumed that a linear trend is synonymous with the photochemical trend. This was a reasonable assumption for the period 1970 through the present, but it will no longer be true for future analyses, as the models predict that the recovery process should now be in its early stages. One analysis question that will now be asked is, "Do we see any evidence for the beginning of the recovery?".
5.4.2 Unresolved longer term variations -- The record that we are using to study long-term variations in ozone is finite in length. The TOMS satellite data goes back only to 1978, though this can be extended back to 1970 if we can put the Nimbus 4 BUV on the same calibration scale. The Dobson network goes back only to the early 1960s and has been intercalibrated to the World Standard Instrument only since 1972. The Arosa Dobson instrument has been measuring since 1928. Sonde measurements of the lower stratosphere go back only to the 1960s.
While these are long records, they do not eliminate the possibility that some longer term variation is contributing to the trend that we see. The satellite ozone record that we use to deduce global trends may well be observing a dynamical component to the trend. The last few northern winters have been exceptionally cold with a stable polar vortex. This makes the last few winters of the record dynamically different from the first few. This could easily show up as a trend in ozone that is not photochemically induced. The true photochemical trend would then be less than the estimate obtained by statistical analyses that do not include this dynamical signal.
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