In the previous sections, we have explored aspects of the vertical structure of the atmosphere, including composition, temperature, pressure, and density with height. We will now explore the horizontal structure of the atmosphere. Differences in the amount of sunlight received at the surface (known as solar insolation) drive the weather and the observed winds on the earth. Temperature changes across Earth's surface (i.e. in the horizontal) are directly linked to the speed and direction of the winds, both at the surface and at different heights. Changes in wind speed in the horizontal and in the vertical direction give rise to instabilities that create weather systems.
A gradient refers to any change over a distance in some physical quantity, such as air pressure, temperature, trace gas concentration, etc. These changes can be in the horizontal or vertical. Temperature changes across the surface are referred to as horizontal temperature gradients. Horizontal temperature gradients and associated changes in windspeed with height are a key aspect of understanding the dynamics of motion that govern the atmosphere.
4.1.1 General characteristics: (i) The troposphere -- In general, tropospheric temperatures decrease uniformly as you go poleward from the equator. In the summer months, the north-south temperature gradient is quite weak until you get at least 40 degrees in latitude (i.e., 40°N and 40°S) away from the equator. Poleward of 40°N/S, the temperature gradient becomes quite tight. This horizontal temperature gradient has an accompanying ribbon of fast moving upper level winds (around 30,000 feet or 9 kilometers in height) known as the jet stream. It is found where the latitudinal (i.e. equator to pole) temperature gradient begins to steepen. The existence of the jet stream is directly linked to the existence of tight latitudinal temperature gradients.
(ii) The stratosphere -- Stratospheric temperatures display a different sort of behavior. As previously stated, the heating due to the presence of ozone causes the temperature to rise with height in the stratosphere. In addition, the horizontal temperature gradient displays a pole-to-pole gradient (as opposed to an equator-to-pole gradient) between the "summer hemisphere" (i.e., the hemisphere where it is summertime) and the "winter hemisphere" (i.e., the hemisphere where it is wintertime). The 6 months of sunlight over the pole of the summer hemisphere heats up the stratosphere here more than anywhere else, while the lack of sunlight over the pole of the winter hemisphere causes the stratosphere here to cool more than any other place. Winds in the stratosphere are strongest at the edge of the "polar night" (i.e., the location between sunlight and 6 months of wintertime darkness inside the Arctic and Antarctic circle). It is here that temperature gradients are strongest. In the summer hemisphere stratosphere, winds weaken and can even reverse directions in the tropics. (This reversal from westerly to easterly winds in the tropical stratosphere is explored in the section 4.2.3 of this chapter and also more thoroughly in Chapter 6.)
4.1.2 Seasonal variation of temperature with latitude --For most of the world's population, large annual changes in weather are a fact of life. These changes are more extreme outside the tropics, where winds and temperature changes can be dramatic. While places in the tropics can be very wet in one season and dry in another, the temperature range is fairly small. It is outside the tropics, in the "extratropics" where the weather changes are most dramatic.
As an example of how temperatures change inside and outside the tropics, the following two figures show temperature changes as a function of altitude and season in two locations. Figures 2.08a and 208b show temperature changes near the equator (2.08a) and 60°N (2.08b) as a function of altitude and season. From these figures we see that in the tropics the amount of solar radiation received does not vary much during the year. The annual changes in temperature, therefore, are fairly small, as demonstrated by the relatively smooth, layered appearance of the isotherms in Figure 2.08a. At high latitudes, however, the tilt of Earth's axis results in significant variations in solar insolation (see Chapter 4). As a result, temperature changes over the course of a year are quite dramatic. Figure 2.08b shows that temperature changes of up to 25 K occur on average in the upper stratosphere and lower troposphere, while smaller changes are found in the middle stratosphere. Naturally, the coldest temperatures are found in winter when the amount of solar radiation incident at high latitudes is at a minimum.
4.1.3 North-south variation of temperature at different heights -- In general, tropospheric temperatures decrease uniformly as you move poleward from the equator while stratospheric temperatures decrease monotomically from the summer pole to the winter pole. Figures 2.09 and 2.10 show the mean temperature distribution at two altitudes, 10 mb (about 30 km) in the stratosphere and 300 mb (about 9 km) in the tropsophere, in the months of January and July, respectively. In both hemispheres temperatures in the troposphere decrease as you move up in height and as you move poleward from the equator. The changes are more severe in the winter hemisphere, with a noticeable drop in temperature poleward of about 30° latitude.
Higher in the atmosphere, the temperature field appears quite different. In the stratosphere, temperatures decrease monotonically as we move from the summer pole toward the winter pole. This is a result of both heating due to ozone and heating due to solar insolation: ozone warms the stratosphere in the winter hemisphere from pole to equator, while solar radiation warms the stratosphere in the summer hemisphere from equator to pole. The combined effect is the monotonic temperature gradient observed. The region of sharpest temperature change appears between 30° and 60° latitude in the winter hemisphere. This places it somewhat poleward of the region of sharpest temperature change in the tropospher.
The structure of temperatures is more zonal during the southern hemisphere winter (July) than is the case in the northern hemisphere winter (January). This is a consequence of two important features: first is the large topographic features in the northern hemisphere, like the Rocky Mountains and the Himalaya-Tibet complex, which are nearly absent in the southern hemisphere; second are the larger land-sea temperature contrasts in the northern hemisphere than those in the southern hemisphere.
In both cases, the coldest temperatures are found over the pole of the winter hemisphere where it is dark for 6 months of the year. This is called the "polar night." A strong jet stream sets up in the lower to middle stratosphere at the edge or "terminator" of the polar night because it is here that solar radiation disappears, allowing temperatures inside the stratospheric polar night to fall to extremely low values, typically in the range of 200K (-100°F) in the northern hemisphere and 190K (-117°F) in the southern hemisphere. Owing to thermal wind balance (see Section 4.1.5), a strong westerly jet stream called the polar night jet sets up. The region of extremely cold temperatures inside the polar night is known as the polar vortex.
4.1.4 Seasonal variation of the zonal wind with latitude -- The zonal wind is the component of the wind field blowing parallel to lines of latitude; i.e., from west to east or vice versa. The zonal jet stream winds discussed here are primarily westerly winds, meaning they blow from west to east. Jet stream winds are westerly whenever temperatures increase from pole to equator. Easterly jet stream winds occur in the stratosphere when the temperature gradient reverses, decreasing from pole to equator. This occurs in the summer hemisphere stratosphere. Meridional winds refer to the north-south winds. When the winds are blowing straight north or south, the zonal wind will be zero.
As we might infer from above, the zonal jets display a significant seasonal variation in those places where the temperature change during the course of the year is significant. Figures 2.11a and 2.11b show the annual variation in the zonal wind as a function of altitude for two different latitudes: the equator (Fig. 2.11a) and 60°N (Fig.2.11b). As was the case with temperature, the zonal wind shows little variability near the equator. In the tropics, the heating provided by solar radiation is fairly uniform both through the year (temporally) and at different locations (geographically). As a result, the temperature field is fairly smooth (as shown above) and the zonal winds fairly weak. Weak temperature gradients and weak zonal winds go together. At high latitudes, however, the zonal wind field changes character more dramatically through the course of the year.
It is around 60°N and 60°S that the greatest changes in zonal winds occur, as it is here that seasonal changes in differential heating between the poles and tropics are maximized. Around 60° latitude, differential heating in the winter hemisphere is much greater than in the summer hemisphere. This results in sharper temperature gradients and hence stronger winter zonal winds than in the summer hemisphere in the vicinity of 60° latitude. This is why the zonal wind here varies so much more during the year than at the Equator, where heating is about constant all year long.
These changes are most notable in the stratosphere, where wind speeds vary by up to 80 m/s (over 100 mph) depending on the season. As we've already mentioned, the most intense winds are found in the winter hemisphere along the edge of the polar night, since it is here that the horizontal temperature gradient is strongest. These winds are the polar night jet.
4.1.5 Variation in the seasonal mean zonal wind with altitude -- We saw in 4.1.3 above that the zonal wind displays a significant seasonal variability in those latitudes where there is a significant change in temperature between the seasons. This is why the variations in zonal wind are much greater at 60° latitude than at the Equator. Figures Fig. 2.11a-b showed how the zonal winds, which are best seen in a plot of wind speed with altitude, are greater at high latitudes where there is greater variability in temperature between the seasons.
In Figures 2.12 and 2.13, we look at how the zonal wind over the globe varies with altitude at different times of the year. Specifically, we look at how the seasonal mean (summer and winter) zonal wind behaves at 300 millibars and 10 millibars. The seasonal mean zonal wind refers to the fact that we have taken an average zonal wind for the month of January (northern winter/southern summer) and July (northern summer/southern winter). These are our seasonal mean zonal winds.
What do these figures show us about the behavior of the zonal wind field in the atmosphere?
The 300 mb level is located near the top of troposphere. It is at this height that we find the strongest zonal winds (blowing from west to east) in the summer hemisphere (coincident with the jet stream) around 45° latitude. This is the location of the subtropical jet stream. In the winter hemisphere, the zonal wind field at 300 mb is stronger than in the summer hemisphere. It also shows more latitudinal variability in the form of more north-south undulations. This variability is related to enhanced wave activity, stronger thermal gradients, and the greater north/south motion of weather fronts that occur in the winter hemisphere.
The 10 mb level is located in the lower stratosphere. One obvious feature is that the zonal wind fields in both summer hemispheres are much smoother. Both summer hemisphere wind fields show a smooth zonal structure with relatively weak winds at all latitudes. Both winter hemisphere wind fields show a much stronger change in zonal wind speeds moving poleward from the equator. This is a consequence of the horizontal (latitudinal) temperature gradient between the equator and pole in the winter hemisphere, especially in the vicinity of the polar night terminator. The polar night jet is found between 40° and 80° latitude.
There are important differences, however, between the winter hemispheres. Note that while the southern hemisphere winter wind field is very zonal, the northern hemisphere winter wind field has a noticeable sinusoidal (wavy) variation around a latitude circle. Also note that as you move poleward in the southern hemisphere winter, the gradient in the zonal wind field is very strong, much stronger than is the case in the northern hemisphere winter. This is a consequence of the colder temperatures found in the winter polar latitudes of the southern hemisphere. The strength of this wind field results in the southern polar vortex being more strongly confined and isolated than is the case with the nothern polar vortex. The isolation of this air mass for an extended period during the winter season plays a large role in the dramatic ozone losses which occur over Antarctica during the southern polar spring (see Chapter 11).
4.2.1 Weather systems and ozone -- The seasonal behavior of the atmosphere has a large effect on ozone levels, but ozone also varies as a result of other atmospheric phenomena. These other phenomena that can be classified by their time scales. Much shorter time scale phenomena involving the day-to-day weather, such as the passage of frontal systems or low pressure areas, can cause 10% variations of ozone just in the course of a few hours. Much longer time scale phenomena, such as the Quasi-Biennial Oscillation (QBO) explained in Section 4.2.3, can result in 10% variations over the course of years.
Tropospheric weather systems are comprised primarily of areas of high and low air pressure. These systems tend to move rather quickly, giving us our daily weather variations. An example of this is provided in the next figure. Figure 2.14 displays images of total ozone (top panels) on December 4 and 5, 1996.
The maximum of ozone on December 4 over North Dakota moves quickly eastward over the following 24-hour period to a position over Illinois. Total ozone over Illinois increases by more than 10% from about 380 Dobson Units to over 428 Dobson Units. The passage of this ozone "high" is associated with a tropospheric weather system shown in the bottom panels. These images are of geopotential height (similar to the pressure charts seen on the nightly weather reports on television).
Figure 2.14 shows us that the ozone "high" is associated with a geopotential "low." The winds move in a counterclockwise sense around these lows. Just as the geopotential low swings eastward, so does the ozone high. The movement of this weather system causes the ozone to move, resulting in the region of high ozone being transported over Illinois. In a similar sense, a high pressure system will lead to a reduction of column ozone. Indeed, so-called "blocking highs" which can persist for weeks can lead to significant reductions in ozone. Why ozone is anticorrelated to geopotential height is a subtle question that we will leave unanswered in this chapter.
Such weather systems as seen in Figure 2.14 are quite common in the extratropics (i.e. outside of the tropical latitudes) of the northern and southern hemispheres. They lead to large day-to-day transport of total column ozone. These weather systems have typical spatial scales of a few thousand kilometers and can influence the overhead amount of ozone for a few days. Thus, these ozone fluctuations are a short-term phenomenon.
4.2.2 Stratospheric sudden warmings -- A second phenomenon that leads to rapid, large ozone changes is the stratospheric sudden warming. As was shown in section 4.1.2, the winter polar stratosphere usually is extremely cold. On an irregular basis, every two to three years, this air mass can rapidly warm up in a very short time period. This warming is illustrated in Figure 2.15. In this figure, we have a plot of the stratospheric mean zonal temperature at 80°N (near the North Pole) at 10hPa or 10 mb. The gray curve is a time-averaged line of zonal mean temperature for 1978-1997. Here the "zonal mean temperature" refers to the average of high and low temperatures at 10mb. The black curve shows temperature fluctuations for July 1989 through June 1990. The stratospheric sudden warming event in question occurs in late January 1990 when zonal mean temperatures warm by about 30 K (or 54 degrees Fahrenheit scale!) in only a few days. Total ozone also changes (rises) substantially in this region over these same few days.
What is the cause of these warmings? They are the result of the rapid displacement of the polar vortex from a roughly symmetric circulation about the pole to a circulation that is offset from the pole. A detailed exploration of this phenomenon, and the more general circulation cell that governs equator to pole stratospheric ozone transport (referred to as the Brewer-Dobson circulation) is given in Chapter 6.
Unlike the January 1990 stratospheric sudden warming event, most such warmings over the Arctic are much more modest. Midwinter warmings of 5-10 K are fairly common in the northern hemisphere. In contrast, mid-winter warmings in the southern hemisphere are quite rare. This is a consequence of the lack of any topography poleward of 50°·S to the edge of the Antarctic continent. This allows frigid air to remain locked up over Antarctica with much less meridional (north-south) mixing (in the form of undulating jet stream wave action). Because the temperature gradient is steeper, the polar night jet is correspondingly stronger. The degree of isolation between stratospheric air inside the polar vortex and outside of it is greater.
The end of the winter season is marked by a rapid warming as sunlight returns to the polar stratosphere and the polar vortex breaks down. Ozone levels and temperatures increase very rapidly. As we shall see in Chapter 11, the presence of manmade chlorofluorocarbon (CFC) compounds and the especially cold conditions in the Antarctic stratosphere combine with the reintroduction of sunlight in the early southern spring (September and October) to destroy large amounts of ozone inside the polar vortex before the vortex breaks down and higher concentration ozone air from lower latitudes intrudes.
4.2.3 The Quasi-Biennial Oscillation (QBO) -- While the polar stratosphere experiences the annual phenomenon of the polar vortex and the periodic stratospheric sudden warming events, the tropical stratosphere experiences its own recurring variability. Specifically, the direction of the winds in the tropical stratosphere are observed to reverse from easterly to westerly and back to easterly again approximately every 26 to 28 months. This reversal of winds in the tropics is known as the quasi-biennial oscillation or QBO.
Figure 2.16 shows a time-versus-height plot of the direction and magnitude of the winds in the tropics as measured by radiosondes launched from Singapore. The altitudes are into the stratosphere at 16 to 32 km. The color bar indicates both the direction and magnitude of the wind. Warm colors (yellow to orange to red) indicate westerly winds with increasing speed. Cool colors (green to blue to violet) indicate easterly winds with increasing speed. Westerly winds are reckoned as positive and easterly winds are reckoned as negative. As we can see from the figure, easterly winds descend about 1 km/month until they reach the tropopause. Just above the descending easterlies are descending westerlies. They follow a similar pattern. Over the course of about 2 years, the winds at any given vertical level in the stratosphere will have gone through a complete cycle, beginning easterly, weakening and reversing to become westerly, weakening and reversing again to become easterly.
The QBO is a regular feature of the tropical stratosphere. Convective activity and other forcings in the tropics generate a variety of atmospheric waves, some of which propagate vertically from the troposphere into the stratosphere (similar to a wave propagating onto a beach). As the waves dissipate in the stratosphere (similar to wave breaking on the shoreline), they deposit their momentum into the stratosphere.
As an analogy, think about water waves at the beach. As waves approach the shoreline, they move relatively freely without much loss of strength. When the wave reaches the shore, the wave height increases due to the decreasing depth of the water below it. As it breaks, it deposits its momentum on the beach, and on any sand castles, pails, and beach balls on the sand. Some of the wave momentum is transferred to the shore, while some is reflected back out to sea. In much the same way, waves moving upward into the tropical stratosphere deposit momentum, and alternately reverse the direction of the background wind field. Effects of momentum deposition on shorelines from water waves are difficult to spot in the short term, while the effects of momentum deposition from tropical atmospheric waves are manifested as the QBO. This momentum, related to the size and propagation speed of the waves, can either be easterly or westerly momentum. Easterly momentum deposition causes an acceleration of the winds, while westerly momentum deposition causes a deceleration of the winds. The alternating wind accelerations and decelerations are related to whether there are easterly winds or westerly winds in the lower stratosphere.
Although the QBO is mainly a tropical phenomenon, its effects are felt well beyond the tropics. Indeed, it turns out that the magnitude of ozone concentrations within the Antarctic polar vortex is related to the phase of the QBO. We will return to stratospheric dynamics in Chapter 6.
4.2.4 The El Niño Southern Oscillation (ENSO) -- Large variations in the ocean temperature of the equatorial Pacific Ocean have become popularly known as the El Niño phenomenon. It is part of a larger, coupled ocean-atmosphere phenomenon known as the El Niño-Southern Oscillation (ENSO). Typically, the waters of the Eastern Pacific near South America are quite cool as a result of upwelling. During a "warm ENSO" or El Niño year, the warmer waters of the Western Pacific migrate eastward, causing a substantial rise in temperatures in the waters of the equatorial eastern Pacific off the coast of Peru and Ecuador. It is this warm water phenomena that was originally referred to as El Niño by Peruvian fishermen who observed its irregular recurrence around Christmas time (hence the name "the Child"). The "Southern Oscillation" part refers to the seesaw of pressure between the Pacific and the Indian Ocean where the normal situation of higher pressures over the tropical eastern Pacific and lower pressures over the tropical central Pacific undergo a reversal.
In order to understand the "oscillating" nature of air pressure between the opposite ends of the tropical Pacific, it is necessary to understand the atmospheric circulation cell that exists along the axis of the tropical Pacific centered on the Equator. Atmospheric pressure is higher over the eastern Pacific and lower over the western Pacific owing to the nature of the "general circulation" of the atmosphere. On the eastern side, there is sinking motion and subsidence, which keeps the climate quite dry. This is enhanced by the rainshadow effects of the Andes mountains and produces the Atacama desert. However, even places like the Galapagos Islands are fairly dry. On the other side of the Pacific, there is rising motion/convergence. Large amounts of water vapor are evaporated in the tropical sun, creating bands of showers and thunderstorms over Indonesia and the western Pacific. The air pressure difference sets up an easterly flow. The flow reverses itself at higher altitudes in the troposphere. This east-west (zonal) circulation cell is referred to as the Walker Cell.
In a warm ENSO event, the Walker Cell weakens and in some cases reverses, so that the normal, strong trade easterlies slacken and even occasionally become westerlies. Rain bands shift eastward. The changes induced by the Southern Oscillation are communicated outside of the tropics as the normal subtropical jetstream is displaced. The opposite situation is called a cold ENSO or La Niña and it occurs when the waters off the coast of Peru and Ecuador become even colder than they usual are. This is associated with a strengthening of the Walker Cell with intensified easterly tradewinds.
4.2.5 Possible correlation between ENSO events and total ozone -- ENSO effects on ozone have been observed in the TOMS total ozone observations [Zerefos et al., 97; Randel and Cobb, 94]. These calculations used a measurement called the Southern Oscillation Index (SOI) to calculate the impact of the ENSO on stratospheric ozone. The SOI is a measurement of difference in sea level pressure (averaged over some period of time) between the island of Tahiti in the central tropical Pacific and Darwin, Australia. The SOI is Tahiti minus Darwin normalized sea level pressure. In the normal or non-ENSO situation, sea level pressure at Tahiti is higher than that at Darwin (just as it is higher at the Galapagos Islands than it is at Tahiti). In a warm ENSO event, the situation reverses, with sea level pressure at Tahiti lower than that at Darwin, so that the SOI has a negative reading. The regression of the SOI against total ozone is shown in Figure 2.17.
There is a weak anticorrelation of the SOI and total ozone in the tropics, with a positive correlation in the southern hemisphere midlatitudes. In general the direct impact of ENSO on stratospheric ozone is quite small, less than 10-20 Dobson Units for relatively strong ENSO events. The exact mechanism by which a strong warm ESNO would reduce stratospheric ozone is unknown.