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.
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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.
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Figure 12-3. Absolute Instability Example Environmental Parcel Temperature Parcel Air Temperature/ Difference Stability Temperature Dewpoint °C 5 0°C 717 -7°C Unstable 4 4°C 9/9 -5°C Unstable 1000's OF FEET 3 Lifting Condensation Level 8°C 11/11 -3°C Unstable (LCL) 2 12°C (14/11.5 -2°C Unstable 1 16°C 17/12 - 1°C Unstable 20°C 20/12.5
12.3.4 Conditional Instability. Conditional instability (see Figure 12-4) is the state of a column of unsaturated air in the atmosphere when its lapse rate of temperature is less than the dry adiabatic lapse rate, but greater than the moist adiabatic lapse rate. An air parcel lifted upward would be initially stable, but at some point above its Lifted Condensation Level (LCL) it would become unstable. The term conditional means the parcel must be lifted to a particular level (altitude) before it becomes unstable and rises because of its own buoyancy. The Level of Free Convection (LFC) is the level at which a parcel of air lifted dry adiabatically until saturated and moist adiabatically thereafter would first become warmer than the surrounding environmental air (i.e., unstable) in a conditionally unstable column of air in the atmosphere. The LFC is a defining feature of a conditionally unstable column of air.
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Figure 12-4. Conditional Instability Example Environmental Parcel Air Temperature/ Temperature Dewpoint °C 5 6°C 717 - 1°C 4 Level of Free Convection 9°C 9/9 0°C (LFC) 1000's OF FEET Lifting Condensation Level 12°C 11/11 JoLt Stable (LCL) 2 15°C (14/11.5 JoLt Stable 1 18°C 17/12 +1°C 20°C 20/12.5 12-5
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12.3.5 Summary of Stability Types. Figure 12-5 below summarizes the possible atmospheric stability types.
Figure 12-5. Stability Types
Temperature Sounding Unsaturated Parcel Saturated Parcel Stability Type A Unstable Unstable Absolute Instability B Neutral Unstable C Stable Unstable Conditional Instability D Stable Neutral E Stable Stable Lapse Absolute Stability F Stable Stable Isothermal G Stable Stable Inversion
12.4 Processes that Change Atmospheric Stability. Changes in atmospheric stability are inversely related to temperature (density) changes with height (see Figure 12-6).
If temperature lapse rates increase, then stability decreases. Conversely, if temperature lapse rates decrease, then stability increases. Most of these changes occur as a result of the movement of air, but diurnal (day/night) temperature variations can play a significant role.
Figure 12-6. Temperature Lapse Rate Effects on Stability Environmental Parcel Temperature Environmental Parcel Temperature Air Temperature/ Difference Air Temperature/ Difference Temperature Dewpoint°C Temperature Dewpoint °C 4 11°C 3/1 +8°C 7°C 8/1 - 1°C 3 12°C 6/1.5 +6°C 10°C 11/1.5 - 1°C 1000's OF FEET 2 13°C 9/2 +4°C 13°C 14/2 -1°C 14ºC 12/2.5 +2°C 16°C 17/2.5 -1°C 15°C 15/3 20°C 20/3 More Stable Less Stable The column of air on the right is less stable because its temperature lapse rate is higher.
12.4.1 Wind Effects on Stability. Wind can act to change the stability of a column of air in the atmosphere by changing the temperature lapse rate. Stability increases when wind blows colder air into the bottom of the air column (cold air advection) and/or warmer air at the top (warm air advection). Conversely, stability decreases when wind blows warmer air into the bottom of the air column and/or colder air at the top.
12.4.2 Vertical Air Motion Effects on Stability. A column of air in the atmosphere will become more stable when it descends (subsides) (see Figure 12-7). As it subsides, it becomes compressed by the weight of the atmosphere and shrinks vertically. The entire layer warms due to adiabatic compression. However, the upper part of the column sinks farther and, thus, warms more than the bottom part. This process acts to decrease the temperature lapse rate and increase stability.
12-7
12-8 Conversely, a column of air in the atmosphere will become less stable when it ascends (rises). As it rises, the rapid decrease in air density aloft causes the column to stretch out vertically. As long as the layer remains unsaturated, the entire layer cools at the dry adiabatic lapse rate (see Table 11-1). However, due to the stretching effect, air at the top of the column cools more than the air at the bottom of the column. This process acts to increase the temperature lapse rate and decrease stability.
Figure 12-7. Vertical Motion Effects on Stability
A rising column of air will become less stable when air at the bottom has a higher relative humidity than air at the top. As the air moves upward, the bottom becomes saturated first and cools at the lesser moist adiabatic lapse rate (see paragraph 11.2). The net effect is to increase the lapse rate within the column and decrease stability. This process is called convective instability, and is associated with the development of thunderstorms.
12.4.3 Diurnal Temperature Variation Effects on Stability. Diurnal (day/night) temperature variations (see Figure 12-8) can have a significant impact on atmospheric stability.
Daytime heating of the surface increases temperature lapse rates and decreases stability.
Conversely, nighttime cooling of the surface decreases temperature lapse rates and increases stability. Diurnal temperature variations are most pronounced in the lower troposphere because air is a poor conductor of heat (see Table 2-3).