10-4 Figure 10-4. Fronts
Fronts are usually detectable at the surface in a number of ways: significant temperature gradients, or differences, exist along fronts (especially on the cold air side); winds usually converge, or come together, at fronts; and pressure typically decreases as a front approaches and increases after it passes.
Fronts do not exist only at the surface of the Earth; they have a vertical structure in which the front slopes over the colder (denser) air mass. Cold fronts have a steep slope, and the warm air is forced upward abruptly (see Figure 10-5). This often leads to a narrow band of showers and thunderstorms along, or just ahead of, the front if the warm rising air is unstable. Warm fronts typically have a gentle slope, so the warm air rising along the frontal surface is gradual (see Figure 10-6). This favors the development of widespread layered or stratiform cloudiness and precipitation along, and ahead of, the front if the warm rising air is stable. Stationary frontal slope can vary, but clouds and precipitation would still form in the warm rising air along the front (see Figure 10-7).
10-5 Figure 10-5. Cold Front
Figure 10-6. Warm Front
10-6 Figure 10-7. Stationary Front
Cold fronts typically move faster than warm fronts, so in time they catch up to warm fronts. As the two fronts merge, an occluded front forms (see Figure 10-8). At the occluded front, the cold air undercuts the retreating cooler air mass associated with the warm front, further lifting the already rising warm air. Clouds and precipitation can occur in the areas of frontal lift along, ahead of, and behind the surface position of an occluded front.
Figure 10-8. Occluded Front
10-7
10.3 The Wave Cyclone Model. A wave cyclone3 is a low pressure circulation that forms and moves along a front. The circulation about the cyclone center tends to produce a wavelike kink along the front. Wave cyclones are the primary weather producers in the mid-latitudes. They are large lows that generally travel from west to east along a front.
They last from a few days to more than a week.
A wave cyclone typically follows a predictable evolution. Initially, there is a stationary front separating warm air from cold air (see Figure 10-9).
Figure 10-9. Wave Cyclone Model—Stage 1
A low pressure wave forms on the front (see Figure 10-10). The front develops a kink where the wave develops. Precipitation develops with the heaviest intensity (dark green) located in the zone of lift along the front.
Figure 10-10. Wave Cyclone Model—Stage 2
As the wave intensifies, both the cold and warm fronts become better organized (see Figure 10-11).
3 A wave cyclone should not be confused with the alternative name for a tornado. They are quite different.
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
10-9
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
11-5
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.