N unit 1, we examined atmospheric physics and how physical properties affect our weather. In this unit, we cover the atmosphere in general. The old saying that "everyone talks about the weather but nobody does anything about it" is becoming increasingly less accurate. As weather journeymen, we can appreciate the advances made in the short history of meteorology as a scientific study. However, we also know the limitations involved in understanding and eventually influencing weather phenomena. Concerning understanding, you might well remember the advice in the book of Proverbs in the Bible: "Wisdom is the principal thing; therefore get wisdom, and with all thy getting (of wisdom), get understanding." This aptly implies that our understanding of any subject depends on the amount of knowledge (or wisdom) at our command. This text material aims toward increasing your understanding of the Earth’s atmosphere.

The first section of this unit covers the atmosphere in general. We review the structure, characteristics of gases, heat balance, and differential heating of the atmosphere. Following that is a discussion of the tropopause and atmospheric circulations. Next comes thorough coverage of jet streams followed by the final section’s topic of air masses. Emphasis on the importance of finding the unique features of each of these topics occurs throughout the section.

The origin of the word atmosphere is Greek, from atmos meaning vapor and sphaira meaning sphere. Literally, it is a sphere of vapor, or a gaseous envelope, surrounding a sphere or globe. The atmosphere of the Earth is an envelope of air surrounding the Earth. Without this envelope, the life forms we consider familiar could not exist. Any planet in the universe may have an atmosphere. However, research has not yet discovered another planet with an atmosphere like that of Earth.

Today, through our increased understanding of the nature of atmospheric phenomena, we can actually alter some weather phenomena. Any ambitious weather modification program, and weather forecasting itself, depends upon our comprehension of the structure of the atmosphere and the interaction of the forces operating within it.

We classify our atmosphere into seven layers and zones. The names of these layers in ascending order are: troposphere, tropopause, stratosphere, stratopause, mesosphere, mesopause, and thermosphere. Figure 2–1 illustrates the relative thickness, height, and temperature of these layers.

The troposphere, whose thickness varies with location and season, has an average height of 38,000 feet (7 miles or 11 km). Within the troposphere, the temperature decreases with height to about –55°C or –60°C. We usually assume a standard temperature decrease in the troposphere of about 6.5°C per 1,000 meters of altitude. Because the troposphere contains the most significant vertical wind motions (unstable) and water vapor and experiences all the weather affecting the Earth’s surface, it captures most of the attention of forecasters.

The tropopause is a thin zone of transition between the troposphere and the stratosphere. Obviously, the tropopause height varies directly with the troposphere thickness and depends largely on geographical location, season, and other factors. For example, over the poles the average tropopause height is only 5 miles, but over the Equator it is 10 miles. Below the tropopause, the temperature normally decreases as height increases. At the tropopause the temperature remains constant with height (isothermal). The tropopause extends only a few thousand feet but acts as the upper limit for the occurrence of weather.

The next layer, the stratosphere, extends from the tropopause to a height of about 154,000 feet (31 miles or 48 km). Temperature within the stratosphere remains isothermal up to 100,000 feet; there it begins to warm with altitude. Temperature reaches a maximum of 7°C (45°F) at the top of the layer.

The stratosphere contains most of the ozone of our atmosphere; it is concentrated between 65,000 and 100,000 feet. The nacreous cloud (mother-of-pearl) also occurs in the stratosphere. Scientists believe there may be an association between the ozone concentration and the nacreous cloud. Because no weather occurs in the stratosphere, flying conditions are excellent.

At the height where the stratospheric temperatures stop increasing, there is a thin zone called the stratopause. It represents a transition between the increasing temperatures of the stratosphere and the decreasing temperatures of the mesosphere.

The mesosphere extends to a height of about 262,000 feet (50 miles). In contrast to the warm upper stratosphere, temperature decreases with altitude in this layer to a minimum near –100°C (–148°F). This rapid temperature decrease causes large convective (vertical) currents. As a result, we consider the mesosphere unstable.

A cloud that occurs in the mesosphere is called the noctilucent cloud (meaning "luminous night clouds") because it appears to glow at night. It is believed that this cloud is composed of meteoric or cosmic and volcanic dust.

Also found within the mesosphere is a concentration of electrons called the "D" layer, which occurs at or about 230,000 feet. This layer reflects radio waves, thus making long distance communications possible.

The last transition zone is the mesopause. This zone marks the end of the temperature decrease within the mesosphere and occurs at or about 262,000 feet. Above this zone extends the atmosphere’s final layer, the thermosphere, which lacks a distinct upper boundary.

Within the thermosphere, the temperature remains isothermal to 300,000 feet and then begins to increase with altitude. This layer contains noctilucent clouds and two layers that reflect radio waves. These layers, like the "D" layer, are concentrations of electrons. They are called the "E" and "F" layers.

An average sample of pure dry air taken from the atmosphere contains (by volume) 78 percent nitrogen, 21 percent oxygen, almost 1 percent argon, and 0.03 percent carbon dioxide. Two other important gases might also be present in variable amounts: ozone (trace amounts) and water vapor (1.2 percent).

This is a colorless, tasteless, odorless, gaseous element. It enters the atmosphere from volcanoes and from the decay of organic matter. Plant life removes nitrogen from the atmosphere.

Oxygen is also colorless, tasteless, and odorless. It is a prerequisite of most forms of animal and plant life. At altitudes above about 13 miles (20 km), the radiation from the Sun breaks the oxygen down into ozone.

Just like nitrogen and oxygen, argon is colorless and odorless. It is considered an inactive gas.

Carbon dioxide, ozone, and water vapor are the more variable gases that significantly affect the weather in the atmosphere.

Carbon dioxide results from the decay of vegetation, combustion, volcanic action, and similar processes. Usually, its maximum concentration occurs around cities and industrial areas. However, the actual amount of carbon dioxide varies with the seasons. There is more combustion and, therefore, more carbon dioxide produced during winter heating.

Ozone reaches a maximum in the stratosphere. It is important because of its ability to "absorb ultraviolet radiation." Without this protection, the Sun would burn everything on Earth.

The third of the variable gases, water vapor, is the most important in determining weather. We could not have the type of weather we are familiar with without water vapor. Water vapor in a parcel of air makes the air lighter than a parcel of dry air. This affects the stability of the atmosphere. The warmer the air, the more water vapor it can hold. The maximum amount of water vapor the air can hold is 4 percent by volume.

Like anything else in the physical world, the Earth’s heat balance is subject to physical processes. These processes have physical laws to describe the relationships involved in radiation and heat exchange. We can use these basic laws to help understand the radiative heat balance between the Earth and the atmosphere. Every phenomenon that we call weather is the result of differential absorption and reradiation of solar energy (discussed earlier) and the availability of moisture.

The Earth’s temperature is decided by the intensity of the solar radiation striking the Earth’s surface. Since the Sun is the basic source of energy for our atmosphere, we must understand how this affects our temperature.

The Sun’s rays strike the Earth’s surface at different angles, which we refer to as the angle of incidence. When this angle is 90° (perpendicular to Earth’s surface), the heat received is the greatest. When the rays are not perpendicular, they must penetrate the atmosphere through a greater distance. This increases the reflection, absorption, and scattering of the rays. When the rays are not perpendicular, less heat is received. The distribution of solar radiation and terrestrial (Earth’s) radiation is an extremely important factor in understanding the weather.

The intensity and wavelength of any emission from a radiating body are dependent on the temperature of the body. The hotter the body, the more intense is the radiation and the shorter are the wavelengths emitted. Since the Sun is the hottest body of our solar system, it emits the shortest wavelengths. Since all objects within our atmosphere are cooler than the Sun, they emit long-wave radiation (discussed earlier).

The significance of the difference in radiation wavelengths lies in the different way they influence the atmosphere. Approximately 35 percent of direct solar radiation is reflected back into space from clouds, the Earth’s surface, and dust in the atmosphere. The Earth absorbs some 51 percent of the incoming solar radiation (insolation) and the atmosphere absorbs the remaining 14 percent.

The portion absorbed by the Earth and the atmosphere must be reradiated to maintain the heat balance within our atmosphere. Since the reradiating bodies are cooler than the Sun, we call their radiation long-wave radiation.

Since incoming radiation heats the atmosphere and the Earth through absorption, these two regions should continue to get hotter and hotter. That this does not happen is due to a complex system of reradiation from the Earth and clouds. Over a long period, the heat loss from the Earth and atmosphere is equal to the heat gain. However, temporary imbalances exist at some points on Earth.

The system is complex because the water vapor in the atmosphere can absorb long-wave (terrestrial) radiation . Thus, the water vapor absorbs and reradiates long-wave radiation to and from the Earth. Because it absorbs long-wave radiation more effectively than short waves, a greenhouse effect is created.

Although the greenhouse effect was briefly discussed earlier, we’ll examine it more in depth to understand this phenomena more clearly. The greenhouse effect is a reference to the effect of clouds and water vapor acting like the glass in a greenhouse when it traps long-wave radiation below it. Here is how it works.

During the day, the Earth absorbs the short-wave radiation that passes through the atmosphere. The heat gained by the Earth is then reradiated back toward outer space as long-wave radiation. When water vapor is present, this long-wave radiation is absorbed by the vapor and then radiated toward the Earth. This process effectively keeps the heat from escaping from that portion of the Earth. This accounts for the higher temperatures on nights when a cloud cover is present.

The water vapor affects the heat balance in one other way that has significance for the meteorologist. When clouds are present during the day, the incoming solar radiation does not reach the Earth’s surface. This is due to reflection of the radiation from the top of the cloud and to some absorption and reradiation outward. About 65 percent of the solar radiation on an overcast day is lost to the Earth through reflection, backscattering, and absorption. This keeps the temperature lower during the daytime in this portion of the Earth.

Earlier, you learned that water vapor is the most important gas in weather formation. Weather also depends on another important variable¾ heat. In fact, how the Earth and its atmosphere are heated is the catalyst for weather formation.

In general, the Earth and its atmosphere maintain a constant average temperature. This is due to a balance between the amount of radiated energy received from the Sun and the amount of radiated energy the Earth loses to outer space. This balance of radiation is maintained by a complex system of reradiation of thermal energy between the Earth and water vapor in the atmosphere.

While it’s true that over long periods the Earth’s temperature is somewhat constant, temperature variations occur because of differential heating of the Earth and its atmosphere. On a global scale, differential heating results from latitudinal variation in the amount of heating of the Earth and its atmosphere. On a smaller scale, other causes of differential heating exist, such as land versus water, hills versus valleys, etc. Differential heating is the true cause of the weather you observe.

The Sun is the source of heat for the Earth. To understand how differential heating occurs, you need to be familiar with factors that influence how the Sun’s energy heats the Earth and its atmosphere. These factors include radiation processes, angle of incidence of solar radiation, absorption, scattering, and albedo.

Radiation from the Sun is responsible for heating the Earth’s surface. Yet, direct solar radiation accounts for only a small part of heat the atmosphere absorbs. Most radiative heating of the atmosphere is the result of reradiated energy from the Earth’s surface. Absorption of radiated energy depends partly on the wavelength of radiated energy. The wavelength of radiated energy depends on the temperature of the emitting body¾ the hotter the body, the shorter the wavelength of emitted energy.

Solar radiation has the full spectrum of electromagnetic energy, including gamma, ultraviolet, visible, infrared, and microwave radiation. Still, as the hottest body in the solar system, the bulk of the Sun’s energy is radiated in short visible wavelengths (less than 3 microns). As a significantly cooler body, the Earth’s radiation is limited to longer infrared wavelengths between 3 and 80 microns.

We can divide the spectrum into two categories¾ short-wave solar radiation (less than 3 microns) and long-wave terrestrial radiation (3 to 80 microns). Figure 2–2 shows the electromagnetic spectrum of radiated energy.

Heating of the Earth and its atmosphere depends on the ability of the individual masses involved to absorb and emit the two types of radiation. Because the atmosphere is largely transparent to incoming short-wave radiation, a large quantity reaches the Earth’s surface where it becomes absorbed and heats the Earth’s surface. The Earth then emits long-wave radiation that is readily absorbed by atmospheric gases, such as carbon dioxide and water vapor.

The only source of short-wave radiation is the Sun. All other mass emits only long-wave radiation. Therefore, any radiation, long-wave or short-wave, absorbed by water vapor and other atmospheric components becomes reradiated as long-wave radiation.

Because of curvature of the Earth’s surface, solar radiation strikes at different angles, called angles of incidence. An angle of incidence is measured from a line perpendicular to the Earth’s surface. When solar radiation strikes perpendicular to the Earth’s surface, the angle of incidence is 0°. If it strikes at angles other than perpendicular, the angle of incidence increases as seen in figure 2–3.

The smaller the angle of incidence, the more the Earth’s surface is heated. The concentration of solar energy is greatest at lower angles of incidence. Notice in figure 2–4 that when the angle of incidence increases, radiation spreads over a larger area. Because the same amount of energy affects a larger area, the Earth’s surface is heated less. This effect becomes magnified over the Earth’s curved surface. There are two separate motions of the Earth that affect angle of incidence¾ rotation and revolution.

The Earth spins about its axis once each 24 hours. Its rotation is responsible for daily heating and cooling as the angle at which solar rays strike the Earth’s surface varies from sunrise to sunset. The angle of incidence is greatest at sunrise and sunset; it is least near noon.

Earth’s revolution

Besides experiencing daily variations in heating, the Earth experiences variations from season to season. These variations are due to the Earth’s annual revolution around the Sun and the tilt of the Earth’s axis. The Earth’s axis is tilted 23.5° from perpendicular to its orbit (fig. 2–5). The tilt is called the angle of inclination. Because of the inclination, as the Earth orbits the Sun, the most direct radiation strikes the Earth at different latitudes throughout the year. The latitude varies between 23.5°N (Tropic of Cancer) and 23.5°S (Tropic of Capricorn).

The seasonal variations in latitude of the Sun’s most direct radiation are marked by four distinct events¾ the summer and winter solstices and the vernal and autumnal equinoxes. June 21 marks the summer solstice when direct solar radiation is over 23.5°N. The winter solstice, December 22, occurs when direct solar radiation is over 23.5°S. The two equinoxes occur when direct solar radiation is over the Equator, 0°. On September 23, the autumnal equinox occurs when the direct solar radiation is shifting from the northern latitudes to the southern latitudes. The vernal equinox (March 21) is just the opposite. Figure 2–6 shows the Earth’s position relative to the Sun for each of these dates.

NOTE: The dates stated for these events are approximate. They vary from year to year because the calendar year is 365 days long. The solar year is 365 1/4 days long.

As we stated earlier, the atmosphere is somewhat transparent to incoming solar radiation. Not all the incoming radiation passes unimpeded through the atmosphere. Only about 53 percent of incoming solar radiation reaches the Earth’s surface. The atmosphere absorbs and scatters some radiation. Also, some surfaces absorb radiated energy better than others. This ability to absorb energy is measured by a surface’s reflectivity or albedo.

Certain gases absorb some wavelengths of incoming solar radiation very well. Most incoming radiation absorbed by the atmosphere is in ultraviolet wavelengths.

The two gases most responsible for absorption of incoming solar radiation are oxygen and ozone. They are so efficient almost no ultraviolet radiation reaches Earth. Atmospheric absorption accounts for a 20-percent reduction of solar radiation reaching Earth’s surface.

Scattering is the deflection of incoming solar radiation in all directions. Scattering of radiation depends to a large extent on particle size. For example, atmospheric particles smaller than the wavelength of visible light (0.5 microns) scatter sunlight. It scatters at a rate inversely proportional to the fourth power of the light’s wavelength. Therefore, shorter wavelengths are much more susceptible to scattering than longer wavelengths.

Atmospheric scattering of solar radiation is apparent in the brightness of the daytime sky. The shorter visible wavelengths of blue and violet light are readily scattered by the particles of the atmosphere. This is why on a clear day the sky is blue. When larger particles are present, such as fog and cloud droplets, all visible wavelengths are scattered equally. This gives the sky the appearance of being white. Consider if the Earth lacked an atmosphere, the sky would be black because there would be no scattering of incoming solar radiation. The atmosphere scatters about 5 percent of the incoming short-wave radiation.

Another important variable controlling the amount of solar radiation absorbed by the Earth is the reflectivity of the Earth’s varied surfaces. Albedo, or the reflectivity of a surface, is the ratio of reflected radiation to average total incoming radiation. Albedo varies greatly from surface to surface and can be as high as 90 percent for fresh snow and as low as 15 percent for bare ground. Using our example from earlier, if you are walking barefoot on a sunny day, you are aware of the difference in a bare patch of ground compared with a cooler, grassy surface. The albedo of the bare ground is 7 to 20 percent while the grass surface’s albedo is 14 to 37 percent. This shows the importance of higher albedo values yielding cooler natural surfaces.

Albedo also explains why the Earth’s surface is cooler on an overcast day. Clouds are also reflecting surfaces with an average albedo between 50 and 80 percent, depending on cloud type and thickness. Clouds in the atmosphere account for another 22 percent of solar radiation that does not reach Earth’s surface. The Earth’s surface reflects another 3 percent of incoming solar radiation.

The tropopause has been defined as a thin zone of transition between the troposphere and the stratosphere. Before World War II, it was believed that the tropopause was a single unbroken layer extending from the Equator to the poles with the height sloping downward toward the poles. However, data gathered during and after the war disproved this idea. The new data shows three or more distinct tropopauses, which occur as leaf-like, overlapping structures. Cells of maximum winds are found among the overlapping leaves. Figure 2–7 illustrates the relationship among the three primary leaves of the tropopause.

The tropopause is one of the most important boundaries in the atmosphere. Below it is the troposphere in which we find all of our weather. The tropopause is of interest to weather journeymen and radar operators who are trying to make determinations about the severity of thunderstorms. This we discuss in more detail later. In this lesson, we look at the characteristics of the tropopause leafs.

Highest of the three leaves that make up the tropopause, the tropical leaf is found near the 100-mb level; the air has a potential temperature (q ) of 100 to 120°C. In winter, this leaf lies along the Gulf Coast of North America. In the summer, it may cover much of the United States and even extend into Canada. Because of its height, the free-air temperature near the tropical leaf is lower than the temperatures at the other leaves.

This leaf occurs at all seasons over the areas suggested by its name. It is found at an intermediate level, usually close to the 200-mb level. The characteristic potential temperature of the air at the mid-latitude leaf ranges from 55 to 75°C.

This leaf usually lies north of the US-Canadian border throughout the year. Occasionally, it drops into the United States during the winter. The polar leaf is the lowest of the three major leaves and occurs around the 300-mb level. The potential temperature of the air is about 30 to 50°C. The polar leaf has smaller free-air temperature extremes than the other leaves. The lower stratosphere above it is comparatively warm.

On mean seasonal cross sections, there appears at all seasons a super high tropical leaf. This leaf is not extensive. It occurs at a level where the pressure is 70 to 80mb, with the potential temperature of the air near 200°C. We find it above the main jet and, therefore, above the areas over which the tropical and polar leaves overlap. Though the persistence of this leaf is recognized, its significance is obscure. The super high tropical leaf is now shown on the National Centers for Environmental Prediction (NCEP) tropopause products. Also, weather journeymen with experience in arctic regions suggest the possibility of an arctic leaf. (Some research has been done on the question of whether an arctic leaf exists.) This leaf is believed to occur where the potential temperature ranges from 10 to 30°C. It is found below the 300-mb level over the United States and Canada during the winter months.

During recent years, radio and television weather reporters have commonly used the term "tropopause." To most people it is still only a word that refers to a level or stratum in the atmosphere. A weather journeyman considers it more important. The temperature discontinuity across the tropopause is not too significant. However, the strong winds resulting from this discontinuity are extremely important to pilots. Strong winds can either cause a pilot extreme costs in time and fuel (headwind) or extend the range and decrease operational time for aircraft flights (tailwind). Additionally, the turbulence associated with the strong wind shear areas can make a flight hazardous.

The tropopause also significantly affects tropospheric phenomena. Research has shown the jet stream is related to surface fronts, cyclones, and severe weather. Because of its lapse rate stability, the tropopause dampens convective activity, thereby reducing or eliminating clouds and turbulence above the level of the tropopause. Strong winds (jets) at the intersection, or overlap, of tropopause leaves have enough intensity to steer surface weather systems and to induce cyclogenesis in the lower levels. The tropopause plays an important part in tornado development. When a particular set of conditions exists, tornado development may follow. This makes it necessary for the weather journeyman to locate the tropopause.

The tropopause is variously described as a boundary, surface, layer, or zone of transition between the troposphere and the stratosphere. Many criteria have been used to define and thereby locate the height or pressure level of the tropopause on a sounding. These criteria are as follows:

• Level of minimum temperature aloft.
• Level of maximum wind speed.
• Level above which there are marked and continuing increases of potential temperature.
• Level at which there is a marked discontinuity in the lapse rate, with the smaller lapse rate above the level.

These location criteria are open to much subjective interpretation. At times, they simply fail to help identify the true tropopause. To lend speed, consistency, and objectivity in locating the tropopause, a rigid, though arbitrary, method has been established. This analysis method takes into consideration that two surfaces (leaves) may overlap.

The predominant (conventional) tropopause is the lowest height (at or between 500 and 30mb) where the lapse rate decreases to 2°C or less per kilometer. Also, it must maintain this average lapse rate for at least 2 km above the point

of stabilization. However, it can occur below 500mb if the following conditions are met:

• It is the only stabilization point that satisfies the definition. Also, the average lapse rate in any higher layer will not exceed 3°C/km over at least a 1 km layer.
• The sounding must reach to at least the 200-mb pressure surface.

Figure 2–8 illustrates a predominant tropopause where the temperature at 400mb (7.2 km) is –50°C and the temperature at 300mb (9.2 km) is –54°C. The vertical distance between the two levels is exactly 2 km, with a temperature difference of 4°C. Therefore, the lapse rate has averaged 2°C per kilometer for at least 2 km above the stabilization point. Thus, it meets the requirement for a predominant tropopause.

When the lapse rate above the predominant tropopause exceeds a 3°C per kilometer average (less stable and getting colder) for at least 1 km, a secondary tropopause exists. This secondary tropopause is at the point where the lapse rate stabilizes, using the same criteria as for the predominant tropopause. In secondary tropopauses, the 1 km interval with an average lapse rate of 3°C/km can occur at any height above the predominant tropopause.

Figure 2–9 illustrates a predominant tropopause and a secondary tropopause. The predominant tropopause is repeated from figure 2–8. Above the predominant tropopause, the temperature decreases 8°C in 1.2 km (6.7°C/km)—greater than 3 °C/km, thus showing a secondary tropopause above. The secondary tropopause occurs at 250mb (10.4 km), where the lapse rate becomes about 0.5°C/km—less than the required 2°C/km.

Useful information can be obtained by analyzing for what we’ll call a significant stabilization point. This is a point where the lapse rate decreases 3°C or more per kilometer. These points of stabilization show where a tropopause leaf is intensifying or weakening, or advancing or retreating. In figure 2–10, the lapse rate from 400 to 350mb is 10°C per kilometer. From 350 to 300mb, the lapse rate has decreased to 6°C per kilometer. The change is 4°C per kilometer, which satisfies the definition of a significant stabilization point at 350mb.

Earlier in this career development course (CDC) volume, you saw how various factors influence heating of the Earth and its atmosphere. For example, angle of incidence is responsible for the wide variation in temperature over different latitudes (differential heating). Because the Equator receives more solar energy than the poles, the equatorial region is warmer. This is not the only factor contributing to horizontal differences. Differences in distribution of land and water masses profoundly affect the overall temperature distribution. From earlier text, we learned that land heats and cools very rapidly, while water heats slowly but retains its heat longer. Extensive snow cover over polar regions reflects much of the incoming solar radiation, further contributing to latitudinal temperature differences.

From the discussion on heat transfer in the last unit, remember convective heating of gases occurs because of currents produced from contact or radiative heating. Convective currents (circulation) due to non-uniform heating of the atmosphere are key factors responsible for the atmospheric circulation pattern. In this section you will study other elements which should help you understand general atmospheric circulation.

You know equatorial areas receive more direct solar radiation than do polar areas. Also, in the lower latitudes, the ratio of land to water is low. Since water loses heat slowly, more heat is received than lost through radiation. In the polar regions, just the opposite is true. The land to water ratio is high and more heat is lost due to radiation than is received. Since the Tropics do not become progressively hotter and polar regions colder, there must be a transfer of heat between the two regions or a general circulation. The mechanism that sets up Earth’s general circulation pattern is the latitudinal transfer of heat.

If the Earth’s surface were smooth, uniform, and stationary, atmospheric circulation would be very simple. The atmosphere would act as a contained fluid and movements within this fluid would be the convective currents caused by temperature and density differences. The latitudinal transfer of heat would result in a single circulation cell.

Unequal heating is the main driving mechanism responsible for the Earth’s atmospheric circulations. The Equator receives more heat than the poles. The warm equatorial air rises and expands until it reaches the tropopause. The tropopause stops the upward motion of this air, causing it to spread and move toward the colder polar regions. Simultaneously, cold and dense air at the poles sinks and starts to flow toward the Equator, replacing the warm air rising in this region. This creates the single cell circulation model possible only on a non-rotating Earth (fig. 2–11).

This theory of circulation best describes the Earth’s general circulation because it considers effects of coriolis force due to the Earth’s rotation. In this circulation model, the Northern and Southern Hemisphere are each divided into three cells of circulation, each spanning 30 degrees of latitude. The latitudes that mark the boundaries of these cells are the Equator, 30° North and South, and 60° North and South. For our purposes, we consider only the Northern Hemispheric cells shown in figure 2–12:

• Hadley.
• Polar.
• Ferrel.

George Hadley, an English meteorologist, theorized this first circulation cell in 1735. The Hadley cell is the strongest of the three cells of circulation and is formed as warm air rises above the Equator and starts to flow northward. The northward flow deflects to the right, due to coriolis, becoming an upper-level westerly flow. As this air moves northeastward toward the pole, it cools and a portion of it sinks at about 30°N. This sinking air spreads northward and southward as it nears the surface. The southward moving air again deflects to the right, becoming the northeasterly trade winds.

Because of the circulation in the Hadley cell, two pressure belts are created. The first is a belt of semipermanent high pressure that results from the sinking air at 30°. This belt of high pressure is called the subtropical ridge. The second pressure belt is a trough of low pressure near the Equator. It is called the near equatorial trough.

This is the northernmost cell of circulation and its mean position is between 60°N and the North Pole. At the pole, cold, dense air descends, causing an area of subsidence and high pressure. As the air sinks, it begins spreading southward. Since the coriolis force is strongest at the poles, the southward moving air deflects sharply to the right. This wind regime is called the surface polar easterlies, although the upper winds are still predominantly from the southwest. Near 60ºN, the southeasterly moving air moving along the surface collides with the weak, northwesterly surface flow that resulted from spreading air at 30°N. This colliding air rises, creating a belt of low pressure near 60°N.

The mid-latitude circulation cell between the Polar cell and the Hadley cell is called the Ferrel cell. This cell is named after William Ferrel, a Nashville school teacher who first proposed its existence. Oddly enough, Mr. Ferrel published his observations in a medical journal in 1856.

The Ferrel cell circulation is not as easily explained as the Hadley and Polar cells. Unlike the other two cells, where the upper and low-level flows are reversed, a generally westerly flow dominates the Ferrel cell at the surface and aloft. It is believed the cell is a forced phenomena, induced by interaction between the other two cells. The stronger downward vertical motion and surface convergence at 30°N coupled with surface convergence and net upward vertical motion at 60°N induces the circulation of the Ferrel cell. This net circulation pattern is greatly upset by the exchange of polar air moving southward and tropical air moving northward. This best explains why the mid-latitudes experience the widest range of weather types.

Differential heating sets into motion and largely sustains the Earth’s circulation pattern. Yet, once in motion, the general circulation pattern is acted upon by a combination of the forces described here.

The pressure gradient force starts the horizontal movement of air over Earth’s surface. It is due to differences in atmospheric density that result in horizontal differences in atmospheric pressure. A pressure gradient is a difference in atmospheric pressure from one point to another. The direction of PGF is always toward lower pressure, perpendicular to the isobars. This is why air starts to flow from high pressure at 30°N toward low pressure at the Equator. The strength of the PGF depends on the horizontal difference in pressure over distance (i.e., pressure gradient). For example, a change in pressure of 20mb over 10 nautical miles produces a greater PGF than a change of 10mb over the same distance. The left diagram on figure 2–13 shows an example of PGF with a change of 10mb over 200 nautical miles. Notice in the right diagram how the PGF is perpendicular to the isobars and points towards lower pressure.

PGF is directly proportional to the strength of the pressure gradient. If the pressure gradient increases (isobars become more closely spaced), then the PGF increases. If the pressure gradient decreases (isobars spread out), then the PGF also decreases.

Contour gradient force (CGF) is the force that represents PGF on a constant-pressure product. It is the rate of height change with change in distance on a constant-pressure surface. Much like hills and valleys on the Earth’s surface, contours correspond to the higher and lower height values, respectively, on a constant-pressure surface.

Differences in virtual temperature and the quantity of mass over a given point in the atmosphere results in changes in thickness between layers. These variations in height result in the contour gradient.

CGF is directly proportional to the strength of the contour gradient. If the contour gradient increases, then the CGF also increases. Conversely, if the contour gradient decreases, then the CGF decreases. Figure 2–14 shows an example of two different contour gradient forces. Notice on the left side of the diagram how the contours are closely spaced (tighter gradient) while the contours spread out on the right side (weaker gradient). The larger arrow indicates stronger CGF on the left while the shorter arrow indicates a weaker CGF on the right. Also notice how both CGF arrows point towards lower heights.

Discussed in detail earlier, this is the force that causes any mass moving free of the Earth’s surface to be deflected from its intended path. The deflection is to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. Deflection due to coriolis is greatest at the poles and is directly proportional to speed of the mass.

Also discussed in detail earlier, this force is the equal and opposite reaction to any center-seeking force. Its strength is directly proportional to speed and radius of rotation.

This force directly opposes and retards the motion of one mass in contact with another. The strength of the force depends on the nature of the contact surface. The more irregular the contact surface, the greater is the frictional force. Therefore, air moving over water is least affected by friction; air moving over mountains is most affected by friction. At heights greater than 3,000 feet above the surface, friction has little effect on air movement. Friction always acts opposite to the direction of motion. With an increase in friction, the wind velocity decreases.

Each force above has an effect on the flow pattern in the Earth’s atmosphere. The two forces having the greatest effect on winds are pressure gradient force and coriolis force. The PGF initiates the horizontal movement of air. As the air starts its motion toward lower pressure, CoF force deflects it to the right. When these two forces are equal, the resulting flow is at right angles to PGF/CGF—this is called geostrophic flow (fig. 2–15). Notice the flow is parallel to lines of equal pressure (isobars/contours) with lower pressure/heights to the left of the direction of flow (V). This is the basis for Buys Ballot’s law that states—"In the Northern Hemisphere, if the wind is at your back, lower pressure is to your left and higher pressure is to your right."

Wind tends to follow the geostrophic pattern (parallel to the isobars) until an imbalance occurs between PGF and coriolis force. Friction is one possible cause for this imbalance to occur. When friction is present, speed is reduced, which lessens the effect of coriolis. Since coriolis is now weaker than PGF, air moves across the isobars toward lower pressure (V_g). Figure 2–16 shows an example of PGF without friction and an example with friction. Friction causes both CoF and CeF to decrease.

Winds tend to back in direction and slow with an increase in friction. For example, if an aircraft were descending from 2,000 feet down to the surface (using the diagram on the left of fig. 2–17), the winds would change from a west-southwest wind to a southwest wind and decrease in speed due to increased friction. Conversely, if the aircraft were ascending from the surface to 2,000 feet, the winds would veer and increase speed.

As we stated before, air tends to flow parallel to lines of equal pressure. This is true even if isobars are curved around high- and low-pressure centers. Curved flow that is parallel to isobars is called gradient flow. Gradient flow, due to its curvature, has the added effect of CeF. To illustrate gradient flow, we examine the balance of forces involved with the two types of pressure circulations.

Air moving around high pressure is termed anticyclonic circulation. In the Northern Hemisphere, anticyclonic circulation is clockwise rotation. To get anticyclonic gradient circulation (V_a), there must be a balance of CoF against PGF and CeF. This occurs if the pressure gradient is constant around the high-pressure center. Figure 2–18 shows the balance of forces for gradient anticyclonic circulation. Also, notice the equal spacing of the circular isobars in the figure. This represents an equal pressure gradient around the pressure center.

If friction affects circulation, air speed is reduced, which again decreases the effects of CoF. Then, as with geostrophic flow, the CoF can no longer balance the PGF and air moves across the isobars toward lower pressure. This is divergent flow—it is diverging or flowing away from the pressure center. Examine figure 2–19 closely to understand the balance of forces associated with divergent anticyclonic circulation.

Low pressure, or cyclonic gradient circulation (V_c), is the exact opposite of high-pressure gradient circulation (V_a), with PGF directed inward. Were it not for CoF, air movement would be directed toward the low center. Still, CoF deflects the movement of air to the right of PGF, resulting in a counterclockwise rotation. As with high-pressure circulation, if we discount friction, the forces balance and a gradient flow circulation results. Notice CoF and CeF now balance against PGF (fig. 2–20).

When we add friction, as we show in figure 2–21, velocity is again reduced, lessening the effect of CoF. The result is the PGF overcomes the CeF and CoF with flow spiraling inward toward the center of the low. This inward directed flow is known as convergent, cyclonic circulation.

Confluence is the merging of wind flow into a common axis of flow (fig. 2–22). As the flow compresses, the air parcels are constricted to a smaller area. Much like putting your thumb over the end of a garden hose, the velocity of the parcels increase.

Difluence is the spreading apart of wind flow from a common axis. The effect is similar to removing your thumb from the end of a garden hose; the velocity of the parcels decrease. (See figure 2–23.)

A supergradient wind (V_super) is a wind which is adjusting to a rapid decrease in CGF downstream. The terms upstream and downstream can be a little confusing; so here’s an example to clarify them. If you were rafting down a river, upstream would be where you came from, while downstream would be where you are going. Now, let’s get back to supergradient winds. The wind temporarily exceeds that of the new, weaker CGF. Supergradient winds occur in areas of difluence where the CGF rapidly decreases. Due to the linear momentum, the air parcels decelerate slowly (temporarily creating an unbalanced situation) and a deflection to the right of flow occurs. (See top diagram figure 2–24.) The bottom diagram of figure 2–24 shows a situation where supergradient winds would occur. Notice how the stronger winds upstream are moving into an area of weaker winds.

CoF causes the deflection because it remains stronger than CGF (or PGF if we were looking at surface features) and remains stronger until the winds slow to the new gradient speed. As the winds deflect towards higher heights, the parcels are converting kinetic energy to potential energy. What do we mean by kinetic and potential energy? First, we must understand the variable relationships of kinetic and potential energy.

Kinetic energy is energy resulting from motion and is directly proportional to mass and velocity. If the mass or velocity of an object is increased, the kinetic energy increases. Conversely, if the mass and velocity of an object are decreased, its kinetic energy decreases.

Webster’s Dictionary defines potential energy as energy that is the result of relative position instead of motion. It is directly proportional to mass and height. If the mass or height of an object are increased, the potential energy increases. Conversely, if the mass and height of an object are decreased, its potential energy decreases.

In our discussion of V_super we stated that parcels are deflected towards higher heights. Therefore, the potential energy increases due to the parcel moving towards the higher heights. We also stated that the winds slow to the new gradient speed. Therefore, kinetic energy decreases due to the velocity decreasing. The process is completed as a conversion and is not an abrupt change from one energy to the other. This is how V_super convert kinetic energy to potential energy.

A subgradient wind (V_sub) is a wind adjusting to a rapid increase in CGF (or PGF) downstream. It is temporarily weaker than that of the new, stronger CGF (or PGF).

Subgradient winds occur in areas of confluence and are in the process of accelerating. Air parcels convert potential energy to kinetic energy by turning towards lower heights or pressure. As potential energy decreases, the kinetic energy increases and the wind crosses contours at slight angles towards lower values. The bottom diagram of figure 2–25 shows slower winds upstream from the stronger winds downstream.

The deflection to the left is caused by the increase in CGF (or PGF). CGF (or PGF) increases immediately upon entering a confluent area and remains stronger than coriolis force until winds increase speed to new gradient speed. While the winds are subgradient, they deflect to the left toward lower heights (or pressure) because of an imbalance of forces between CGF (or PGF) and CoF.

Cyclostrophic wind (V_cyc) is a wind that occurs when CGF (or PGF) is balanced by CeF. CoF is negligible when compared to CGF (or PGF) or CeF. This circulation can only exist around low pressure areas where CGF (or PGF) is exactly opposite CeF. (See figure 2–26.)

This circulation cannot form around a high-pressure area because CGF (or PGF) and CeF both point in the same direction. (See figure 2–27.)

Some examples of cyclostrophic wind flow in the atmosphere include hurricanes, tornadoes, dust devils, and waterspouts. Hurricanes occur in the Tropics where CoF is small.

In the mid-latitudes, some systems with a very strong PGF and a small radius of rotation result in a large CeF, often forming dust devils, tornadoes, or waterspouts. Due to the extreme PGF (or CGF), either clockwise or counterclockwise flow is possible around these small-scale, low-pressure systems. CoF is only a very small part of the balance of forces since CeF and PGF (or CGF) are very large.

The actual wind is difficult to observe accurately because of the many disruptions to the flow (gust, eddies, turbulence, vertical motion, etc.) and our observation limitations.

The distribution of land and water over the surface of the Earth alters the general circulation pattern. The unequal distribution of land and water masses amplifies the effects of unequal heating of the Earth’s surface. In addition, the belts of high and low pressure described earlier are actually broken into separate cells of high and low pressure. Besides these semipermanent features, a system of migratory pressure systems contributes to the general atmospheric circulation system.

Semipermanent pressure systems are largely the result of the general circulation pattern. They are called semipermanent because they can usually be found at the same place with the same intensity during the same month each year. You should become familiar with several of these semipermanent high- and low-pressure systems.

The belts of high pressure described as existing near 30° in both hemispheres are actually divided into distinct circulation centers. These centers comprise the subtropical ridge. The locations of the centers of these large anticyclonic circulations are over the oceans between the continents. In the Northern Hemisphere there are two: (1) the North Pacific high, located over the Pacific Ocean to the west of California and (2) the North Atlantic or Azores high, located west of the African coast near the Azores. A third high-pressure circulation exists southeast of the United States near Bermuda. This circulation, an extension of the North Atlantic high, is called the Bermuda high. The locations of the three Southern Hemispheric subtropical highs are in the South Pacific, South Atlantic, and Indian Oceans.

Subtropical high-pressure systems exist throughout the year, though seasonal variations in location and intensity do occur. Circulation centers migrate north and south with the seasonal variation in the solar angle of incidence. The maximum intensity of these systems occurs in the summer months for their respective hemispheres. For example, the North Atlantic system is strongest in July, while the South Atlantic system is strongest in January.

There is another persistent type of high-pressure system in the Northern Hemisphere—the Siberian high. Unlike the subtropical systems, this high-pressure circulation system exists only during the winter. It is a low-level feature that forms because of the cold, dense air present in this region during the winter. This pressure system has its mean position at 50°N and is over central Asia.

Another common persistent high-pressure circulation is the Polar high. This high-pressure circulation is also the result of cold temperatures. Though its name suggests a center over the North Pole, its center varies with the seasons—usually displaced toward the area of coldest temperatures. During the winter months, it appears as an extension of the Siberian high. In the summer, it is situated near the cooler air of the North Pacific. This high is also called the Greenland or Canadian high.

An area of low pressure is along the Equator. It is called the near-equatorial trough. It appears as an elongated area of lower pressure circling the globe. It exists due to strong equatorial warming and convergence of the northeasterly trade winds of the Northern Hemisphere and the southeasterly trade winds of the Southern Hemisphere. Its position varies with the seasonal change in latitude of direct solar radiation. The near-equatorial trough moves northward during the Northern Hemisphere summer and southward during the Northern Hemisphere winter.

There are three other semipermanent low-pressure circulations important to general circulation in the Northern Hemisphere—the Icelandic, Aleutian, and Asiatic lows. Two lows corresponding to the semipermanent belt of low pressure at 60°N are the Icelandic and Aleutian lows.

The Icelandic low is prevalent throughout the year. It is located in the North Atlantic near Iceland and Greenland. Its existence is due to cold temperatures of the ice caps of Iceland and Greenland contrasting to warmer water temperatures of the North Atlantic Ocean. This contrast is greatest in the winter; therefore, the Icelandic low is most intense then.

The location of the Aleutian low is in the Gulf of Alaska near the Aleutian Islands. Its presence is explained by a contrast in land and water temperatures, as was the Icelandic low. Yet, the Aleutian low is less persistent in the summer because of the influence of the Polar and North Pacific highs.

The Asiatic low dominates the entire Asian continent during the summer. The mean position of the center of this low is east of the Persian Gulf and just north of 23°. In this place, the Asiatic low dominates the same region as the Siberian high but in different seasons. The Asiatic low forms because increased summer heating in the Southern Asian Deserts replaces the cold Siberian high that dominates Asia in winter.

Semipermanent low pressure in the Southern Hemisphere is located in a somewhat continuous band around the globe at 60°S. This is very close to the pressure pattern suggested by the three-cell theory. This continuous belt exists because of a lack of significant land masses at this latitude in the Southern Hemisphere. In contrast, the Northern Hemisphere has a large concentration of land area in the latitudes around 60°.

The other areas of prevalent low pressure in the Southern Hemisphere are over the continents during January. These low-pressure circulations are the result of rapid heating of these land masses during the summer of the Southern Hemisphere.

Changes in the circulation pattern discussed thus far have been caused by seasonal changes in semipermanent pressure systems. Yet, the general circulation pattern of the atmosphere includes wind systems that continually migrate. Migratory systems comprise both anticyclonic (high-pressure) and cyclonic (low-pressure) circulations. They are responsible for the rapidly changing weather conditions found especially in the middle latitudes. This occurs because the migratory systems are responsible for the transportation of polar and tropical air into the middle latitude regions. When these two types of air with greatly different characteristics clash, the result is stormy weather, such as violent spring thunderstorms or heavy winter snowfalls. Usually, fair weather is associated with migratory high-pressure systems while stormy weather is associated with migratory lows.

By dictionary definition, tertiary circulations are circulations of a third order. In the atmospheric circulation system, tertiary circulations are small-scale wind systems that occur with the general circulation pattern. They are a result of the Earth’s rough surface and temperature differences between land and water.

Tertiary circulations are frequently called local winds and have names that link them to the place where they occur. In spite of their many names, we can lump tertiary circulations into four categories—local-cooling, local-heating, adjacent heating/cooling, and forced circulation winds.

Local-cooling winds are the result of differing specific heats of land surfaces in a generally small place. These are nighttime circulations that develop as one surface radiates heat energy and cools faster than another, such as the mountain breeze shown in figure 2–28. At night, the rocky slopes of the mountain lose their heat more rapidly than the moist, vegetation-covered surfaces in the valley. Air in contact with the mountain cools and becomes denser than the valley air.

This density difference causes a pressure gradient to form with high pressure on the mountains and low pressure in the valley. Winds then blow in the direction of PGF, down the slopes into the valley. Wind speeds associated with this type of wind rarely exceed 15 knots, even with a well-developed breeze.

Local-heating winds are the result of differential heating of two surfaces. Since it is due to local heating, it is a daytime feature dependent on incoming solar radiation. The reverse of the mountain breeze, the valley breeze, illustrates this type of wind (fig. 2–29). In the valley breeze situation, the Sun’s rays strike the mountain slopes before they strike the valley surfaces. Thus, the mountains warm quicker than the valley and lower pressure forms on the mountains. Now the PGF is directed up the slopes and the breeze blows from the valley to the mountain. The valley breeze is usually a more developed wind than the mountain breeze and reaches its maximum intensity in the mid-afternoon.

Land and sea breezes illustrate the adjacent heating/cooling winds shown in figure 2–30. Since land masses absorb and radiate heat three times faster than water masses, land is warmer in the daytime and cooler at night. Differential heating of the land and water causes lower pressure over the warmer land and higher pressure over cooler water. Therefore, PGF is directed toward the land from the water. Thus, on a coast or shoreline, sea breezes (coming from the water) would be expected in the afternoon. At night, the system reverses itself as the land cools more rapidly than the water. Therefore, higher pressure exists over land than the water and the PGF is directed offshore. The best-developed land breeze occurs just before dawn.

The entire pattern depends on the fact that land and water heat and cool at different rates. The sea breeze reaches maximum velocity between 1400 and 1600 local time and is much stronger than the land breeze. Sea breezes may be felt as far inland as 25 or 30 miles. The sea breeze is also a major factor in afternoon thunderstorm development along coastal regions. Cool air blowing inland creates a boundary (sea-breeze front) between cool, moist marine air and warm inland air. The sea-breeze front is often seen in Florida during the summer. Land and sea breezes are also common along rivers and lake shores.

Forced-circulation winds are circulations that intensify because of terrain, usually mountain ranges. There are two types of winds in this category—the fall wind and the foehn.

Fall winds (sometimes called glacier winds) are cold downslope winds that remain colder than the air they replace in the lowlands. Fall winds are extremely strong and gusty (60 to 100 knots) and often cause considerable damage. These winds begin in the highlands where air is trapped over cold, snow-covered regions until forced downslope by a migratory pressure system. The winds move down the mountain slope and intensify as they move through the valleys. They are most violent in winter and spring, with single occurrences lasting for several days. Fall winds are frequently called Mistrals, the name of the local wind in the lower Rhone valley of southern France. Another name for a fall wind is the Bora that blows from the mountains of Yugoslavia to the Adriatic sea.

The foehn wind, called the chinook in North America, is also a downslope wind, but foehn winds are warm and dry. The drying results when air is forced up the windward side of a mountain. This rising air cools and loses its moisture due to condensation. It then continues over the crest of the mountain and down the slopes, warming adiabatically as it compresses in the lower elevations. The result is a warm, dry wind that blows downward along the leeward slopes of the mountain. A foehn wind may raise the temperature as much as 50°F in the low lands on the leeward side of a mountain. The temperature increase may occur in as little as a few minutes. Figure 2–31 shows the sequence of events for a foehn wind and the clouds that you may observe on the windward side and near the peaks of the mountain. Notice the absence of cloudiness on the leeward side of the mountain. Expect this since the air loses its moisture through condensation as it is forced up the windward side.

Figure 2–31. Foehn (chinook) wind.

Though considered a local wind, the foehn wind can be a hazard to aviation. This is the case when the wind direction is perpendicular to the mountain and the wind speed increases rapidly with height. This causes extremely turbulent conditions above the peaks and leeward of the mountains. This condition may be detected by the presence of cap, rotor, and lenticular clouds as depicted in figure 2–32. Also notice in the figure the wind currents associated with mountain-wave conditions.

Identify each of these described pressure systems as semipermanent or migratory and as a high or low type of pressure system.

The discovery of the jet stream during the closing days of World War II introduced insights on the general atmospheric circulation. Jet streams now are recognized as potent forces for inducing cyclogenesis, tornadoes, squall lines, mountain waves and turbulence. Regions where jet streams are not present tend to have more uniform and less violent weather.

The jet stream awareness also improved our understanding of the mid-latitude climatic structure and enabled forecasters to strengthen the reliability of both short-range and long-range prediction procedures.

We can best describe the basic structure of a jet stream as a river flowing horizontally through the atmosphere. They are normally thousands of miles in length, hundreds of miles wide, and a few miles deep. Wind speeds in a jet stream vary along each of its dimensions. Looking into the spiraling column in the direction of flow, you can see the change in speed across the jet stream width and through its height (fig. 2–33). The greatest wind speeds through the jet stream are in the core within the interior of the jet stream. Notice the displacement of the core to the left and top within the jet stream. Therefore, change in wind speed over distance (speed shear) is greatest above and to the left of the jet stream core as you look in the direction of flow.

The core of the jet stream, found along its length, defines the jet stream axis (fig. 2–34). Notice, the axis is drawn along the band of maximum winds when we view it from the top. The figure also shows the variation in wind speed along the length of a jet stream.

Jet maxima (maxes) or jet streaks are intermittent regions of the strongest wind speeds within the jet. The isotachs which define the jet maximum are elliptically (football) shaped, both horizontally and vertically. (See figure 2–35.) Core speeds of 100 to 150 knots are common within jet maxima, but speeds up to 300 knots have been observed. Jet maxima are not stationary features but really move along the jet stream axis. The importance of jet maxima comes from their link with other atmospheric disturbances.

The jet stream is a narrow belt of strong winds in the upper troposphere with speeds of 50 to 200 knots. In the Northern Hemisphere, these winds usually have a westerly component. Jet stream position varies between different latitudes and elevations around the Earth. It even varies in latitude and elevation within a small geographical area. The jet stream may appear as a continuous band around the Earth, but more often it gradually diminishes at one or more points and then reappears farther downstream.

The jet stream is significantly wider than it is deep. Though the core may be more than 100 miles wide, it may be at least 10,000 feet thick. This produces stronger vertical wind shears than horizontal shears.

Winds in the jet core average 120 to 150 knots during the winter, with frequent occurrences above 200 knots. Since the increase to these high speeds occurs in a comparatively short distance, the shear values are also high. The highest vertical shear usually occurs above 500mb.

Shear above the core is usually greater than the shear below the core because on the average the wind gradient is greater above the core. The wind usually decreases 50 percent within 5 km above the core, but takes 5.5 km to decrease 50 percent below the core.

Cross sections of the jet stream also indicate horizontal wind shear. There is a rapid decrease of wind speed on each side of the jet core. This decrease can be as much as 100 knots in 100 miles on the north side and 100 knots in 300 miles on the south side. This indicates greater horizontal shear and turbulence north (cold air side) of the jet. (See figure 2–36.)

A jet stream rarely occurs alone in the atmosphere; multiple jets are the rule. Two major jet streams are frequently observed: the polar front jet and the subtropical jet. Figure 2–37 shows the average position of these two Northern Hemisphere jet stream cores with respect to the tropopauses. The jet core depicted between the polar and mid-latitude tropopauses is called the polar front or mid-latitude jet.

The core between the mid-latitude and tropical tropopauses is called the subtropical jet.

A third jet stream, called the arctic or polar night jet stream, is believed to occur during the winter at 75°N around the 300- to 350-mb level near the Arctic circle. It is most often observed in the winter, hence the name polar night. In addition to these three jets, the tropical easterly jet is located over southern Asia and northern Africa. This jet is most frequently observed in the summer at the 150-mb level. In the following discussion, we focus our attention on the two major jet streams affecting the Northern Hemisphere—the polar and subtropical jet streams.

Jet streams are primarily caused by two reasons. One origin of a jet stream is caused by a large horizontal temperature contrast associated with a large thermal wind vector. Jet streams that form by this large temperature contrast have a jet axis that is at the tropopause because this is where the contour gradient force is at a maximum. Above and below the tropopause, the slope of the constant-pressure surface weakens. Horizontal temperature contrast is the primary originator of the polar front jet stream.

The second origin of a jet stream is due to the conservation of angular momentum. Newton’s first law of motion states that an object at rest tends to stay at rest and an object in motion tends to stay in motion at the same velocity (be conserved) until it is acted on by an outside force.

Angular momentum is the product of the mass times velocity and radius of rotation for an object, or parcel, following a curved path. It is stated by the formula:

Angular momentum is conserved for an object when forces acting upon the body are in balance. If the mass is held constant and the radius decreases, then the velocity must increase in order to conserve angular momentum. This is known as an inversely proportional relationship. Our earlier example pertaining to a ball on the end of the string may help explain this. If you have a ball on the end of a string and you spin it around your finger, the string begins to wrap around your finger. As the string wraps around your finger, the length of string decreases (decreasing the radius) and, ultimately, the ball spins faster (increasing its velocity).

Because the atmosphere is a fluid and not attached to the solid Earth, it tries to conserve its angular momentum. As a parcel moves poleward, the radius (of the Earth) decreases with an increase in latitude, causing its velocity to increase. Conversely, with a parcel flowing equatorward, the radius increases with a decrease in latitude; as a result, its velocity decreases.

The PFJ is the primary jet stream in the middle latitudes of the Northern Hemisphere. It may appear as a continuous band around the globe, but, more often, it vanishes at one point and reappears downstream. Its disappearance and reappearance is often due to variations in its height. Its existence can be attributed to large horizontal temperature differences (thermal gradient) across the upper and middle latitudes. These temperature differences exist not only at the surface, but vertically to the tropopause. The stronger the thermal gradient, the stronger is the jet stream.

Vertically the PFJ is at the top of the troposphere near the tropopause zone. In fact, the polar jet divides the tropopause into two segments called leaves (fig. 2–38). The PFJ is within the warm air, south of the coldest air. It frequently marks the boundary between cold polar air and warmer subtropical air. In the summer, with temperature differences lessened by warming in the middle and higher latitudes, the PFJ is found at higher latitudes. When the middle and higher latitudes cool in winter, the polar jet migrates farther south. Winter is also the time of greatest horizontal temperature differences between latitudes and the time when the PFJ’s speeds are greatest.

The PFJ follows a generally eastward track; still, if you view the entire pattern outlined around the globe, you should notice large north to south undulations. The undulations, called long waves, consist of large low-pressure troughs and high-pressure ridges. The series of waves the jet stream follows is the long-wave pattern. Observations reveal the series of long waves created by this flow number as few as three and as many as seven around the globe—the average number found is five (fig 2–39).

Jet maxima in the PFJ are of prime importance to the meteorologist because of their link to mid-latitude migratory pressure systems. PFJ maxima cause and steer the major mid-latitude tropospheric pressure systems. Low-pressure systems are found ahead of the maxima while high-pressure systems move behind them.

A valuable clue in locating the PFJ is the relationship between the jet stream and the polar front. The core of the PFJ is directly above the isotherm ribbon associated with the frontal zone (thermal concentration) on the 500-mb surface. This relationship has a high degree of correlation, since troughs and ridges slope very little between 500 and 300mb and have almost no slope from 300 to 200mb.

There are several methods of locating the 500-mb frontal zone. The simplest is to locate the position of the –17°C isotherm. The polar front usually cuts the 500-mb surface along the –17°C isotherm. The concentration of isotherms are located to the warm side of the transition zone, with the largest concentration between –17 and –20°C.

The other methods usually depend on locating areas of highly concentrated isotherms. A jet stream is usually associated with a thermal concentration of more than 10°C in 200 miles. Warm air is south of the concentration and cold air is north. Thus, with your back to the wind (facing downstream), the cold air is to your left and the warm air to your right.

Since the temperatures reverse above the tropopause, in a north-south direction, the jet stream will be just below this temperature reversal. This reverse occurs above 200mb. Thus, the jet stream core is usually located between 300 and 200-mb levels.

From these observations come the following rules:

• The 500-mb position of the jet stream is located in the –17 to –20°C isotherm ribbon, with the most frequent occurrence along the –17°C isotherm.
• The jet stream core has been found between 8,960 meters (29,400 ft) and 9,236 meters (30,300 ft) more than 82 percent of the time during the winter.
• The width of the jet core is about equal to the width of the 500-mb zone having the maximum temperature concentration (isotherm ribbon).

Another jet stream important to the Northern Hemisphere is the subtropical jet (STJ) stream. It exists independently, but frequently exists with the PFJ.

Two main factors contribute to the formation of the STJ. The first is the convergence of the Hadley and Ferrel Cells as previously discussed in atmospheric circulation. The second factor is the horizontal temperature contrast due to the interaction of warm, moist subtropical air and cold dry polar air.

The STJ position is south of the polar front jet and is normally observed between 25 and 30°N, with a mean latitude of 28° N. It extends from the Hawaiian Islands eastward to southern Florida. Conservation of angular momentum is a factor that accelerates the subtropical jet, while coriolis force causes recurvature to the right. Remember, as the jet moves northward, radius decreases and velocity increases. At the same time, coriolis force is also increasing as it moves closer to the poles, thereby increasing curvature.

We can trace the STJ’s axis around the entire globe. In contrast to the polar jet stream, the STJ persists at the 200- to 150-mb level. Though the STJ persists at higher altitudes than the polar jet stream, STJ winds as high as 120 knots have been observed as low as 300mb, but rarely below this level. The STJ reaches its maximum speed at its farthest projection poleward. Reference figure 2–37, which shows the relative vertical and horizontal positions of the STJ and PFJ’s. Notice the STJ also divides the tropopause into leaves with the tropical leaf to the south and the mid-latitude leaf to the north.

The STJ typically flows more directly from west to east. At times, the subtropical jet moves northward and merges with the polar front jet but is more often observed as a separate feature. However, when these two jet streams do interact, favorable conditions for extreme turbulence exist. This condition also favors the development of severe thunderstorms, as satellite imagery often confirms severe weather activity as a characteristic wedged-shaped cloud pattern in this region. The STJ is very important in transporting warm moist air poleward to interact with cool dry polar air. This produces strong temperature, contour, and isotach gradients associated with severe weather.

It is in the region of confluence of the polar and subtropical jet streams where the STJ velocities are the greatest. This occurs partly because the STJ axis is at its furthest poleward extent and, consequently, the radius of the Earth is smaller (conservation of angular momentum). Also, the confluence of the STJ with air already moving at high speeds in the PFJ results in producing a broader and stronger jet stream.

The STJ is less connected with lower tropospheric pressure systems than the PFJ. Occasionally, it does become superimposed on a lower tropospheric pressure system and can exert considerable influence on the pressure system’s development. Just like the PFJ, the STJ also has seasonal variances, being stronger in the winter than in the summer, and typically weak during hurricane season.

Jet streams undergo a distinct life cycle consisting of a period of organization and one of disintegration. Their horizontal structure, as it appears on upper-level products, depends on the phase of this life cycle. When the jet is well organized, the longitudinal axis follows the pattern of long waves. Also, its amplitude is comparable to that of the 300-mb contours. Thus, the jet stream can provide a clue of the long-wave pattern.

Figure 2–40 shows a 300-mb product with the contours removed to emphasize the horizontal structure of the jet stream. A lack of uniformity in wind speed along the jet stream axis is clear in the isotach analysis. As shown, the wind gradient along the axis is sharp and the wind speed difference between the maximum and minimum approaches 100 knots. Often, the difference between individual maximum centers exceeds 100 knots.

During disorganized periods, the well-defined wind speed concentrations of 50 knots or more tend to disappear. When this occurs, the westerlies either break into closed circulation patterns or show comparatively uniform velocities throughout the middle latitudes. The westerly wind speed pattern during periods of dissipation is characterized by the development of many, comparatively weak "jet fingers," often no more than 300 to 400 statute miles apart. Figure 2–41 shows a 300-mb product with "jet fingers."

The jet stream follows the contours closely. Small departures are often noted, however, when the wind speed increases suddenly in a region of comparatively weak contour gradient or decreases suddenly in an area of strong contour gradient.

Diverging contours downstream cause the jet to deflect toward greater heights (higher pressure at the same altitude) and converging contours downstream tend to turn the jet toward lower heights.

The isotach maxima that occur along the axis of the jet stream are related to the short-wave troughs in the westerlies and their connected surface systems. The maxima usually move near the speed of these waves (average 35 to 40 knots). Distances between maxima vary from about 10 to 25 degrees longitude and speeds may reach 250 knots.

The maxima tend to be closed, symmetrical, lenticular-shaped isotachs that keep their shape over short periods. They undergo definite life cycles with periods of formation and decay, which often may be followed on synoptic products. Figure 2–42 illustrates the life cycle of a fully organized jet stream. Beginning with a balanced symmetrical arrangement of isotachs along the jet stream axis (fig. 2–42, view A), the isotachs downstream from the maxima centers are displaced faster than those behind the center (fig. 2–42, view B). In addition, high-speed isotachs are displaced faster downstream than slower ones. This increases the area enclosed by the highest valued isotach and packs the isotachs downstream from the maxima centers (fig. 2–43, view C). The maxima merge with the low-speed area downstream and the low-speed area disappears (fig. 2–42, view D). The resulting pattern leads back to the beginning stage (fig. 2–42, A), and the cycle is repeated. For a jet stream on the decline, the condition shown in figure 2–42, view C, is followed by that in figure 2–42, view E, rather than 2–42, view D. This stage is the first indication of the disintegration of the jet stream and its breakup into jet fingers.

Occasionally, the jet stream surrounds the Earth as a continuous band, but usually diminishes gradually at one or more points. It then reappears in a new location. This new appearance is often connected with migration of the jet stream.

The jet stream can migrate due to changing orientation at speeds near 10 knots, but it does not, at any time, move perpendicular to its axis. This movement usually is to the north or south.

Basically, the jet stream follows the migration of the Sun. Its mean position advances south in winter and retreats toward the poles during summer. The average latitude of the jet stream in the Northern Hemisphere is 35 to 40°N in winter and 42°N in summer.

In a jet migrating north, a west wind maximum usually emerges from the Tropics and gradually moves through the lower mid-latitudes. Another maximum, initially located in the upper mid-latitudes, advances toward the Arctic Circle and weakens. The jet stream is usually well-organized and the contour pattern shows troughs extending into low latitudes.

As the jet progresses northward, the amplitude of the long waves decreases and cut-off lows, south of the prevailing westerlies, dissipate.

The jet stream moves southward until the long-wave amplitude reaches a critical limit. At this stage, the jet usually disintegrates. Also, at this point, cut-off, closed lows frequently form in the southern portion of the troughs. This reestablishes a zonal flow now farther to the north and a new jet stream forms to start another cycle.

Cyclones and bad weather tend to be associated with jet streams. This is especially true for weather associated with a warm front in the formative and mature stages of a cyclone.

You learned before that the mid-latitude jet usually lies directly above the 500-mb position of the polar front. Using average frontal slopes, this places the jet stream about 300 miles behind the surface cold front and about 600 miles ahead of the surface warm or stationary front.

For instance, the following rules apply:

• The jet stream will be perpendicular to an occlusion and to a north-south oriented cold front with no associated warm front.
• The jet stream will remain north of an unoccluded wave cyclone.
• The jet stream will be south (near the point of occlusion) of the low associated with an occluded front.
• In a series of lows of a cyclone family, each low will be associated with a jet stream maximum.
• Each cyclone associated with westerly flow aloft will have an associated isotach maximum.
• The jet stream will parallel the direction of the warm sector isobars of a surface low.
• The jet stream will roughly parallel the isobars around the southern periphery of a cold (slow-moving) surface low.
• The jet stream will roughly parallel the isobars around the northern periphery of a warm (slow-moving) surface high.
• When a cold, moving, polar high stagnates and begins to warm up, the jet stream usually changes markedly. As the thermal contrast in the southwest quadrant is destroyed and the warm-air advection moves northward, the original jet stream usually dissipates and a new jet forms to the north.

Four cloud patterns are usually associated with the jet stream:

• Lines of cirrus in bands.
• Patches of cirrocumulus or altocumulus castellanus.
• Waves of altocumulus.
• Lenticular clouds in waves.

These cloud patterns often extend from horizon to horizon and move very rapidly. Clouds occur most frequently in the following positions:

• 10,000 to 15,000 feet below the jet core and 4 to 5° poleward.
• 5,000 to 10,000 feet below the jet core and 4 and 5° equatorward.

By far, most clouds are on the equatorward side of the jet. Most of the time, clouds are not observed above the jet core.

The greatest occurrence of precipitation usually straddles the jet axis, with a slight bias toward the poleward side.

Air motions within the atmosphere are normally very complex. However, sometimes a large portion of the atmosphere comes to rest (stagnates) or moves slowly over a uniform surface. When this occurs, the air within this portion of the atmosphere assumes the same thermal and moisture properties as the surface beneath it. These properties tend to be uniform over the entire mass of air.

When the mass of air is set in motion, it maintains these characteristics until it is modified by the surfaces over which it moves. Therefore, you must know where these air masses form, how they are formed, how they are modified, and the flying conditions they produce.

Three factors are considered necessary for the formation of air masses. First, there must be a surface with comparatively uniform properties, especially the properties of temperature and moisture. This surface may be water or land. Second, the air must stagnate over the uniform surface. Third, there must be a large divergent flow. This flow inhibits temperature contrasts and produces uniform characteristics throughout the air mass.

The stagnation of air over a surface of uniform conditions permits a condition of temperature equilibrium between the ground and air, actually the formation of the air mass. This is accomplished by:

• Turbulent-convective transport of heat upward into the higher levels of the air.
• Cooling of air by radiation loss of heat.
• Transport of heat by evaporation and condensation processes.

By far the most effective and fastest transfer is the turbulent-convective transport of heat upwards. Least efficient and slowest is radiation. The equilibrium layer created by the turbulent-convective process has the greatest vertical depth, but layers resulting from radiation cooling are comparatively shallow. Evaporation and condensation contribute in each of the other two processes, but are very ineffective in creating air masses by themselves.

Anticyclonic systems are excellent for air-mass formation. They are usually stagnant or slow moving, which allows the air time to adjust its heat and moisture content to that of the underlying surface. They have a divergent airflow, which spreads the properties horizontally over a large area, with turbulence and convection distributing these properties vertically. These factors have the effect of creating uniformity throughout the air mass.

Two examples of the anticyclonic systems are the warm highs, Bermuda and Pacific, and the Siberian cold high. Warm highs extend to great heights, due to higher thickness aloft, and produce an air mass of comparatively large vertical extent. Cold highs have a moderate or shallow vertical extent and thereby produce air masses of moderate or shallow height.

Cyclonic systems are not conducive to air-mass formation. They are characterized by strong wind speeds that prevent the air mass from remaining over a surface long enough to become modified. Cyclonic systems also have a convergent windflow that creates temperature contrasts. The convergence creates ascending air, which causes the air to lose contact with the surface. Without this contact, the air cannot acquire the properties of the surface.

The world’s belts of convergence resemble the cyclonic systems. They also are not conducive to air-mass formation.

These characteristics decide the types of air masses forming in various areas of the world (source regions). Tropic regions are covered predominantly by water surfaces, thus providing the uniform surface.

Air masses in the polar regions form slowly through the loss of heat by radiation. The surface of the polar region is almost completely covered by snow, thus providing the needed uniform surface.

The temperate latitudes provide neither of the two factors needed for air-mass formation and rarely serve as source regions. Steep gradients of temperature and a large variance in the land/water distribution prevail in temperate latitudes. Instead of the temperate zone being a source region, the temperate zone is a transition zone. Air masses with different properties move in, clash, and produce a variety of weather.

There are two notable exceptions where air masses form despite the lack of proper factors. These are the air masses that form maritime polar air over the North Pacific between Siberia and North America and over the Atlantic water off Labrador and Newfoundland. Both areas have active and strong windflow, yet the air masses still form. This formation, however, is related to the characteristics of the region rather than the system.

The Glossary of Meteorology defines an air mass as a widespread body of air having properties identified by its horizontal composition—particularly temperature and humidity. Also, vertical temperature and moisture variations are about the same throughout the horizontal extent of a given air mass. Air masses should not be classified as a part of the secondary circulation, but air-mass properties, particularly along the boundary zone between contrasting air masses, cause circulation phenomena that are secondary in scale. This is the way we examine air-mass properties.

If atmospheric circulation was strictly zonal, we would find only two different air masses—polar and tropical. The boundary, or zone, between them is the polar front. This distribution is clear-cut in the middle and upper troposphere. At the surface, perturbations in the zonal flow confuse the picture. Horizontal turbulence, variations in insolation, and the motion of the air masses themselves invoke complicated transitional effects.

The considerable differences in the earth-atmosphere parameters influence the different sections of the atmosphere. Radiation rates, convection, evaporation, condensation, insolation, and turbulence are all functions of the Earth’s surface over which a given air mass may spend a large part of its time. Air masses over oceans acquire maritime characteristics: mild temperatures, considerable moisture (at least in the lower layers), moderate lapse rates, and a degree of conditional instability. Polar air masses tend to be very cold and dry with temperature inversions in the lower levels and fairly uniform lapse rates aloft. These examples serve to show that we should examine some source regions of the principal air masses. This enables us to better understand their characteristics.

We classify air masses with letter identifiers according to their source region and characteristics. The source region is considered the most useful criterion. Thus, the primary air-mass identifier refers to the source region. A capital P represents a polar air mass and tropical air masses are represented by a capital T. On occasions, arctic air-mass properties are important; a capital A identifies these air masses. Equatorial air masses are represented by a capital E. To further classify the source region, we make a distinction between land and water areas since these regions give decidedly different characteristics to overlying air masses. Use the letter "c" for continental or "m" for maritime surfaces; for example, cP indicates a continental polar air mass and mT a maritime tropical air mass.

The stability characteristics of air masses often depend on the temperature differences between them and the surfaces over which they may travel. The third category identifies whether the air is colder (k) or warmer (w) than the surface it is moving over. Therefore, "w" designates those that are warmer and "k" those that are colder than the surfaces over which they are found. This classification is relative and frequently difficult to determine. The several possibilities can be listed thus:

We can explain the system as follows. mPw represents a polar maritime air mass that is warmer than the surface over which it moves. cPk describes a continental polar air mass that is colder than the surface over which it moves, and so on. We do not need all the categories for arctic air masses. They are usually continental in character and colder than the surface over which they move.

Continental and maritime designations signify the influence of the surface on air-mass properties. Arctic, polar, and tropical primary designations suggest the importance of latitude of the source regions. Typical k air masses have convective clouds, turbulent gusty winds, and good visibility. Heating of these air masses, by a warmer surface, increases the lapse rate and vertical eddies that favor convective transport. On the other hand, w air masses have their lower layers cooled by the colder surface. This tendency toward an inversion of the lapse rate increases stability. When present cloud systems are stratiform and vertical turbulence is suppressed, visibility is frequently poor with low-level radiation and advection fog the rule.

A word of caution is in order, however. The interaction of air masses with the surface is not the only mechanism that produces observed weather effects. Convergence, divergence, and subsidence are very important in finding lapse rates and stability. Dynamical processes frequently dominate the thermal effects.

We use a fourth letter to represent the stability of the air mass. An "s" signifies a stable air mass and "u" specifies an unstable air mass.

By combining these letters, we can obtain a relatively complete description of the air masses. A classification of cPws signifies a stable, continental polar air mass that is warmer than the surface it is moving over.

Figures 2–43 and 2–44, respectively, show January and July air-mass source regions. We’ll limit our emphasis to those air masses and source regions that influence the Northern Hemisphere weather.

These have their source regions in central Canada and Siberia. Next to the extreme Arctic and Antarctic regions, these are the coldest places on Earth. This is because of the intense radiational cooling and lack of insolational heating in these source regions. The surface is completely frozen and is mostly snow and ice covered. Air masses in these regions are extremely cold and stable and are, therefore, very dry. Clouds are uncommon in the cP air masses, although ice fogs frequently form when temperatures fall below –40°F.

These have their source regions in the central portion of high-latitude continents. Insolational heating warms the surface substantially, even to the point of dissipating the snow cover and thawing the ground to a noticeable depth. Summertime cP air masses are cool and dry, but not necessarily stable. They are fundamentally changed forms of the wintertime cP air masses that have experienced insolational heating in their lower layers. Generally, their lapse rates are less steep than their wintertime counterparts.

We find these air masses over open oceans at high latitudes. The cool, moist air masses indigenous to this locale are mainly cP air masses that have moved out over the open ocean. The lower layers are substantially changed by the comparatively warmer, yet still cold, water surface. The mP air masses tend to be moist and unstable in the lower layers, and cold and dry aloft.

These start in basically the same regions as the wintertime mP air masses. For the North American continent, these regions are the Gulf of Alaska and the North Atlantic. Summertime mP air masses resemble their wintertime counterpart; they are cool and moist in the lower layers, and cool and dry aloft. The overall temperature is somewhat higher than in the winter mP air masses. Instability is more general in the lower layers and abrupt change in the lapse rates is found in the moisture discontinuity aloft. An inversion at this level is not uncommon.

We find these air masses over the open oceans, just south of the polar front, near 30°N latitude. Indeed, mT air-mass source regions extend southward across the near-eq