10-8 Figure 10-11. Wave Cyclone Model—Stage 3
In the fourth stage, the wave becomes a mature low (see Figure 10-12). The occluded front forms as the cold front overtakes the warm front.
Figure 10-12. Wave Cyclone Model—Stage 4
As the cold front continues advancing on the warm front, the occlusion increases and eventually cuts off the supply of warm moist air (see Figure 10-13). This causes the low to gradually dissipate.
Figure 10-13. Wave Cyclone Model—Stage 5
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10.4 Dryline. A dryline is a low-level boundary hundreds of miles long separating moist and dry air masses. In the United States, it typically lies north-south across the southern and central High Plains during the spring and early summer, where it separates moist (mT) air from the Gulf of Mexico to the east and dry desert (cT) air from the southwestern states to the west (see Figure 10-14).
Figure 10-14. Dryline Example
The dryline typically advances eastward during the afternoon and retreats westward at night. However, a strong wave cyclone can sweep the dryline eastward into the Mississippi Valley, or even further east, regardless of the time of day. Low-level clouds and early morning fog often prevail in the moist air, while generally clear skies mark the dry side. Severe and sometimes tornadic thunderstorms often develop along a dryline or in the moist air just to the east of it, especially when it begins moving eastward.
A typical dryline passage results in a sharp drop in humidity (hence the name), clearing skies, and a wind shift from south or southeasterly to west or southwesterly. Blowing dust and rising temperatures also may follow, especially if the dryline passes during the daytime. These changes occur in reverse order when the dryline retreats westward.
11-1 CHAPTER 11. VERTICAL MOTION AND CLOUD FORMATION
11.1 Introduction. A cloud is a visible aggregate of minute water droplets and/or ice particles in the atmosphere above the Earth’s surface. Fog differs from cloud only in that the base of fog is at the Earth’s surface while clouds are above the surface.
Clouds form in the atmosphere as a result of condensation of water vapor in rising currents of air, or by the evaporation of the lowest layer of fog. Rising currents of air are necessary for the formation of vertically deep clouds capable of producing precipitation heavier than light intensity.
11.2 Vertical Motion Effects on an Unsaturated Air Parcel. As a bubble or parcel of air ascends (rises), it moves into an area of lower pressure (pressure decreases with height).
As this occurs, the parcel expands. This requires energy, or work, which takes heat away from the parcel, so the air cools as it rises (see Figure 11-1). This is called an adiabatic process. The term adiabatic means that no heat transfer occurs into, or out of, the parcel.
Air has low thermal conductivity (see Table 2-3), so transfer of heat by conduction is negligibly small.
The rate at which the parcel cools as it is lifted is called the lapse rate. The lapse rate of a rising, unsaturated parcel (air with relative humidity less than 100 percent) is approximately 3 °C per 1,000 feet (9.8 °C per kilometer). This is called the dry adiabatic lapse rate. This means for each 1,000-foot increase in elevation, the parcel’s temperature decreases by 3 °C. Concurrently, the dewpoint decreases approximately
0.5 °C per 1,000 feet (1.8 °C per kilometer). The parcel’s temperature-dewpoint spread decreases, while its relative humidity increases.
This process is reversible if the parcel remains unsaturated and, thus, does not lose any water vapor. A descending (subsiding) air parcel compresses as it moves into an area of higher pressure. The atmosphere surrounding the parcel does work on the parcel, and energy is added to the compressed parcel, which warms it. Thus, the temperature of a descending air parcel increases approximately 3 °C per 1,000 feet
(9.8 °C per kilometer). Concurrently, the dewpoint increases approximately
0.5 °C per 1,000 feet (1.8 °C per kilometer). The parcel’s temperature-dewpoint spread increases, while its relative humidity decreases.
11-2 Figure 11-1. Unsaturated Ascending/Descending Air Parcel Example
At the surface, the air parcel has a temperature of 18 °C and a dewpoint of 3 °C, which indicates that it is unsaturated. As the parcel ascends, its temperature decreases at the dry adiabatic lapse rate of 3 °C per 1,000 feet, while the dewpoint decreases at
0.5 °C per 1,000 feet. The temperature-dewpoint spread decreases while relative humidity increases. The process is reversible if the parcel remains unsaturated.
11.3 Vertical Motion Effects on a Saturated Air Parcel. The Lifting Condensation Level (LCL) is the level at which a parcel of moist air lifted dry adiabatically becomes saturated. At this altitude, the temperature-dewpoint spread is zero and relative humidity is 100 percent.
Further lifting of the saturated parcel results in condensation, cloud formation, and latent heat release. Because the heat added during condensation offsets some of the cooling due to expansion, the parcel now cools at the moist adiabatic lapse rate, which varies between approximately 1.2 °C per 1,000 feet (4 °C per kilometer) for very warm saturated parcels to 3 °C per 1,000 feet (9.8 °C per kilometer) for very cold saturated parcels.
Concurrently, the parcel’s dewpoint decreases at an identical rate. For simplicity, examples shown in this AC use a moist adiabatic lapse rate of 2 °C per 1,000 feet.
Regardless of temperature, the relative humidity remains constant at about 100 percent.
11-3 As the saturated air parcel (see Figure 11-2) expands and cools, however, its water vapor content decreases. This occurs because some of the water vapor is condensed to water droplets or deposited into ice crystals to form a cloud. This process is triggered by the presence of microscopic cloud condensation (and ice) nuclei, such as dust, clay, soot, sulfate, and sea salt particles. The cloud grows vertically deeper as the parcel continues to rise.
Figure 11-2. Ascending Air Parcel that Becomes Saturated Example
At the surface, the air parcel has a temperature of 18 °C and a dewpoint of 13 °C, which indicates that it is unsaturated. As the parcel ascends, its temperature decreases at the dry adiabatic lapse rate of 3 °C per 1,000 feet, while the dewpoint decreases at 0.5 °C per 1,000 feet. The temperature-dewpoint spread decreases while relative humidity increases until the parcel achieves saturation at its LCL of 2,000 feet. As the parcel continues to ascend, condensation produces cloud formation. Because the heat added during condensation offsets some of the cooling due to expansion, the parcel now cools at the moist adiabatic lapse rate of 2 °C per 1,000 feet. The parcel’s dewpoint decreases at an identical rate as the lost water vapor condenses to form the cloud.
The relative humidity of the ascending saturated (i.e., cloudy) parcel remains constant at about 100 percent.
A descending saturated air parcel (see Figure 11-3) quickly becomes unsaturated.
Its temperature increases 3 °C per 1,000 feet, while its dewpoint increases at 0.5 °C per 1,000 feet (see Table 11-1). The temperature-dewpoint spread increases while relative humidity decreases.
11-4 Figure 11-3. Descending Air Parcel Example
At 5,000 feet, both the temperature and dewpoint of the air parcel are 6 °C, which indicates that it is saturated. As the parcel descends, it quickly becomes unsaturated.
Its temperature increases 3 °C per 1,000 feet, while its dewpoint increases at 0.5 °C per 1,000 feet. The temperature-dewpoint spread increases while relative humidity decreases until the parcel reaches the surface. Note that the parcel is now much warmer and drier at the surface then when it began the vertical motion process in Figure 11-2.
Table 11-1. Air Parcel Vertical Motion Characteristics
PARCEL UNSATURATED SATURATED Temperature Change Dewpoint Change Relative Humidity Temperature Change Dewpoint Change Relative Humidity Ascending (rising) -3 °C/1,000 ft -0.5 °C/1,000 ft Increases -1.2 °C to -3 °C/1,000 ft Identical to temperature change 100% Descending (subsiding) +3 °C/1,000 ft +0.5 °C/1,000 ft Decreases
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11.4 Common Sources of Vertical Motion. There are many sources of vertical motion in the atmosphere. Four of the most common types of vertical motion are orographic effects, frictional effects, frontal lift, and buoyancy.
11.4.1 Orographic Effects. Winds blowing across mountains and valleys cause the moving air to alternately ascend and descend. If relief is sufficiently great, the resulting expansional cooling and compressional warming of air affects the development and dissipation of clouds and precipitation.
For example, a mountain range that is oriented perpendicular to the prevailing wind flow forms a barrier that results in a cloudier and wetter climate on one side of the range than on the other side (see Figure 11-4). As air is forced to rise along the windward slope, it expands and cools, which increases its relative humidity. With sufficient cooling, clouds and precipitation develop at and above the LCL. Conversely, on the mountain’s leeward slope, air descends and warms, which reduces its relative humidity, and tends to dissipate clouds and precipitation. In this way, mountain ranges induce two contrasting climatic zones: a moist climate on the windward slope and a dry climate on the leeward slope. Dry conditions often extend hundreds of miles to the lee of a prominent mountain range in a region known as rain shadow.
Figure 11-4. Orographic Effects Example
The air parcel begins with a temperature of 15 °C, dewpoint of 10 °C, and a relative humidity of 80 percent at 2,000 feet. As the parcel is lifted on the windward slope, the temperature cools at the dry adiabatic lapse rate of 3 °C per 1,000 feet and the dewpoint
11-6 cools at a rate of 0.5 °C per 1,000 feet until it becomes saturated at the LCL at 4,000 feet.
Then, the air parcel’s temperature and dewpoint both cool at the moist adiabatic lapse rate of 2 °C per 1,000 feet until the parcel reaches the summit at 12,000 feet. At that altitude, the parcel’s temperature is -7 °C, the dewpoint is 7 °C, and the relative humidity is 100 percent. As the air parcel descends the leeward slope, the temperature increases at a rate of 3 °C per 1,000 feet while the dewpoint increases 0.5 °C per 1,000 feet. The air parcel ends with a temperature of 23 °C, dewpoint of -2 °C, and a relative humidity of 33 percent at 2,000 feet, much warmer and drier than at the beginning.
Orographic effects are especially apparent from west to east across the Pacific Northwest, where the north-south Cascade Range intercepts the prevailing flow of humid air from the Pacific Ocean. Exceptionally cloudy, rainy weather prevails western slopes, whereas semiarid weather characterizes the eastern slopes and areas further east.
11.4.2 Frictional Effects. In the Northern Hemisphere, the surface wind spirals clockwise and outward from high pressure, and counterclockwise and inward into low pressure due to frictional force. The end result is that winds diverge away from surface high pressure, causing the air to sink, compress, and warm, which favors the dissipation of clouds and precipitation. Conversely, winds converge into surface low pressure, causing the air to rise, expand, and cool, which favors the formation of clouds and precipitation given sufficient moisture (see Figure 11-5).
Figure 11-5. Frictional Effects
11.4.3 Frontal Lift. Frontal lift (see Figure 11-6) occurs when the cold, denser air wedges under the warm, less dense air, plowing it upward, and/or the warmer air rides up and over the colder air in a process called overrunning. Cloud and precipitation will form given sufficient lift and moisture content of the warm air.
11-7 Figure 11-6. Frontal Lift
11.4.4 Buoyancy. Air near the ground can warm at different rates depending on the insular properties of the ground with which it is in contact. A newly plowed field will warm faster than an adjacent lake. These temperature differences result in different densities, allowing the warm air to become buoyant. The denser cool air will tend to push (i.e., lift) the less dense warm air aloft. On a grand scale, the tendency of air to rise due to heating, and how high it will rise, is referred to as stability and is covered in Chapter 12.
12-1 CHAPTER 12. ATMOSPHERIC STABILITY
12.1 Introduction. Convective clouds and precipitation pose a distinctly different flying environment than stratiform clouds and precipitation. These sharply contrasting conditions result from the atmosphere either resisting or accelerating the vertical motion of air parcels. Atmospheric stability is the property of the ambient air that either enhances or suppresses vertical motion of air parcels and determines which type of clouds and precipitation a pilot will encounter.
12.2 Using a Parcel as a Tool to Evaluate Stability. An air parcel can be used as a tool to evaluate atmospheric stability within a specified vertical column of air in the atmosphere. A parcel is selected from a specified altitude (usually the surface) and hypothetically lifted upward to a specified test altitude. As the parcel is lifted, its temperature decreases due to the expansion and latent heat effects discussed in Chapter 11.
The parcel and the surrounding environmental air temperatures are then compared. If the lifted parcel is colder than the surrounding air, it will be denser (heavier) and sink back to its original level. In this case, the parcel is stable because it resists upward displacement. If the lifted parcel is the same temperature as the surrounding air, it will be the same density and remain at the same level. In this case, the parcel is neutrally stable. If the lifted parcel is warmer and, therefore, less dense (lighter) than the surrounding air, it will continue to rise on its own until it reaches the same temperature as its environment. This final case is an example of an unstable parcel. Greater temperature differences result in greater rates of vertical motion.
12.3 Stability Types. The stability of a column of air in the atmosphere is classified by the distribution of parcel stabilities within the column. The Earth’s surface is typically selected as the base while the top determines the column’s depth. Five unique types of atmospheric stability can be identified.
12.3.1 Absolute Stability. Absolute stability (see Figure 12-1) is the state of a column of air in the atmosphere when its lapse rate of temperature is less than the moist adiabatic lapse rate. This includes both isothermal and inversion temperature profiles. An air parcel lifted upward would be colder (denser) than the surrounding environmental air and would tend to sink back to its level of origin.
Figure 12-1. Absolute Stability Example Environmental Parcel Temperature Parcel Air Temperature/ Difference Stability Temperature Dewpoint °C 5 18°C 717 + 11°C Stable 4 19°C 9/9 +10°C Stable 1000's OF FEET 3 Lifting Condensation Level 20°C 11/11 +9°C Stable (LCL) 2 21°C (14/11.5 +7°C Stable 1 20°C 17/12 +3°C Stable 20°C R0/12.5
12.3.2 Neutral Stability. Neutral stability (see Figure 12-2) is the state of a column of air in the atmosphere in which an ascending (or descending) air parcel always has the same temperature (density) as the surrounding environmental air. If the column of air is unsaturated, then neutral stability exists when its lapse rate of temperature equals the dry adiabatic lapse rate. If the column of air is saturated, then neutral stability exists when its lapse rate of temperature equals the moist adiabatic lapse rate.
12-2
8/23/16 Figure 12-2. Neutral Stability Example Environmental Parcel Air Temperature/ Temperature Dewpoint °C 5 7°C 717 0°C 4 9°C 9/9 200 1000's OF FEET 3 Lifting Condensation Level 11°C 11/11 0°C Neutral (LCL) 2 14°C /14/11.5 0°C Neutral 1 17°C 17/12 0°C Neutral 20°C 20/12.5
12.3.3 Absolute Instability. Absolute instability (see Figure 12-3) is the state of a column of air in the atmosphere when it has a superadiabatic lapse rate of temperature (i.e., greater than the dry adiabatic lapse rate). An air parcel displaced vertically would be accelerated in the direction of the displacement. The kinetic energy of the parcel would consequently increase with increasing distance from its level of origin.
12-3