AFH15-101 Ch. 1 Surface Weather Elements PDF

AFH15-101 5 NOVEMBER 2019 9 Chapter 1 SURFACE WEATHER ELEMENTS 1.1. Visibility. The American Meteorological Society defines visibility as “The greatest distance in a given direction at which it is just possible to see and identify with the unaided eye, (1) in the daytime, a prominent dark object aga...

AFH15-101 5 NOVEMBER 2019 9 Chapter 1 SURFACE WEATHER ELEMENTS 1.1. Visibility. The American Meteorological Society defines visibility as “The greatest distance in a given direction at which it is just possible to see and identify with the unaided eye, (1) in the daytime, a prominent dark object against the sky at the horizon, or (2) at night, a known, preferably unfocused, moderately intense light source.” Forecasting visibility is a challenge due to the difficulty in predicting the complicated behavior of “dry” and “moist” (both liquid and solid) airborne particles that obstruct or reduce visual range. Basic overviews of dry obstructions (lithometeors) and moist obstructions (hydrometeors) are provided below, followed by more concentrated analyses of individual obstructions. 1.1.1. Dry Obstructions (Lithometeors). A lithometeor is the general term for particles suspended in a dry atmosphere; these include haze, smoke, dust, and sand. 1.1.1.1. Haze. Haze is an accumulation of very fine dust or salt particles in the atmosphere; it does not block light, but instead causes light rays to scatter. Haze particles produce a bluish color when viewed against a dark background, but look yellowish when viewed against a lighter background. This light-scattering phenomenon (called Mie scattering) also causes the visual ranges within a uniformly dense layer of haze to vary depending on whether the observer is looking into the sun or away from it. Typically, haze occurs under a stable atmospheric layer and significantly effects visibility. In general, industrial areas and coastal areas are most conducive to haze formation. 1.1.1.2. Smoke. Smoke is usually more localized than other visibility restrictions. Accurate visibility forecasts depend on detailed knowledge of the local terrain, surface wind patterns, and smoke sources (including schedules of operation of smoke generating activities). 1.1.1.3. Blowing Dust and Sand. Windblown particles such as blowing dust and sand can cause serious local restrictions to visibility, often reducing visibility to near zero. The critical wind speed for lifting dust and sand varies according to vegetation, soil type, and soil moisture. Specific forecasting rules vary by station and time of year; local references and procedures should document the wind speeds, directions, and surface moisture conditions in which visibility restrictions are most likely to occur. 1.1.2. Moist Obstructions (Hydrometeors). Condensation or sublimation of atmospheric water vapor produces a hydrometeor. These particles form either in the free atmosphere or at the earth‘s surface, and include frozen water lifted by the wind. Hydrometeors that cause surface visibility reductions generally fall into two categories: 1.1.2.1. Precipitation. Precipitation refers to all forms of water particles, both liquid and solid, which fall from the atmosphere and reach the ground; these include: liquid precipitation (drizzle and rain), freezing precipitation (freezing drizzle and freezing rain), and solid (frozen) precipitation (ice pellets, hail, snow, snow pellets, snow grains, and ice crystals). 10 AFH15-101 5 NOVEMBER 2019 1.1.2.2. Suspended (Liquid or Solid) Water Particles. Liquid or solid water particles that form and remain suspended in the air (damp haze, cloud, fog, ice fog, and mist), as well as liquid or solid water particles that are lifted by the wind from the earth‘s surface (drifting snow, blowing snow, blowing spray) cause restrictions to visibility. One of the more unusual causes of reduced visibility due to suspended water/ice particles is whiteout, while the most common cause is fog. 1.1.2.2.1. Whiteout Conditions. Whiteout is a visibility-restricting phenomenon that occurs when a uniformly overcast layer of clouds overlies a snow- or ice-covered surface. Most whiteouts occur when the cloud deck is relatively low and the sun angle is at about 20° above the horizon. Cloud layers break up and diffuse parallel rays from the sun so that they strike the snow surface from many angles (Figure 1.1). This diffused light reflects back and forth between the snow and clouds until the amount of light coming through the clouds equals the amount reflected off the snow, completely eliminating shadows. The result is a loss of depth perception and an inability to distinguish the boundary between the ground and the sky (i.e., there is no horizon). Low-level flights and landings in these conditions become very dangerous. Figure 1.1. Whiteout Conditions. 1.1.2.2.2. Fog. Fog is often described as “a stratus cloud resting near the ground.” Fog forms when the temperature and dew point of the air approach the same value, either through cooling of the air (producing advection, radiation, steam, or upslope fog) or by adding enough moisture to raise the dew point (frontal fog). When composed of ice crystals, it is called ice fog. 1.1.2.2.2.1. Advection Fog. Advection Fog forms due to moist air moving over a colder surface, and the resulting cooling of the near-surface air to below its dewpoint temperature. Advection fog can occur over both water and land. 1.1.2.2.2.2. Radiation Fog. Also called ground or valley fog, this type of fog is produced by radiational cooling. Under stable nighttime conditions, long-wave radiation is emitted by, and cools, the ground, forming a temperature inversion. In turn, moist air near the ground cools to its dew point. Depending on the moisture content of the ground, moisture may evaporate into the air, raising the dew point of this stable layer, accelerating radiation fog formation. AFH15-101 5 NOVEMBER 2019 11 1.1.2.2.2.3. Upslope Fog. Upslope fog occurs when sloping terrain lifts air, cooling it adiabatically to its dew point and saturation. Upslope fog may be viewed as either a stratus cloud or fog, depending on the point of reference of the observer. Upslope fog generally forms at higher elevations and builds downward into valleys. This fog can maintain itself at higher wind speeds because of increased lift and adiabatic cooling. Upslope winds more than 10 to 12 knots usually result in stratus rather than fog. The eastern slope of the Rocky Mountains is a prime location for this type of fog. 1.1.2.2.2.4. Frontal Fog. Associated with frontal zones and frontal passages, this type of fog can be divided into types: warm-front pre-frontal fog; cold-front postfrontal fog; and front-passage fog. Pre- and post-frontal fogs are caused by rain falling into cold stable air and raising the dew point. Frontal-passage fog can occur in a number of situations, such as when warm and cold air masses, each near saturation, are mixed by very light winds in the frontal zone. It can occur when relatively warm air is suddenly cooled over moist ground with the passage of a well-marked precipitation cold front. It can also occur during low-latitude summer, where evaporation of front-passage rain water cools the surface and overlying air enough to add sufficient moisture to form fog. 1.1.2.2.2.5. Ice Fog. Ice fog is composed of ice crystals rather than water droplets and forms in extremely cold, arctic air (–29°C (–20°F) and colder). Factors contributing to reduced visibility associated with ice fog are temperature, time of day, water vapor availability, and pollutants. Burning hydrocarbon fuels, steam vents, motor vehicle exhausts, and jet exhausts are major sources of water vapor and pollutants that help to produce ice fog. A strong low-level inversion contributes to ice fog formation by trappingand concentrating the moisture in a shallow layer. Once ice fog forms, it usually persists until the temperature rises or there is an air mass change. 1.1.2.2.2.6. Sublimation Fog. The Glossary of Meteorology defines sublimation as the “transition from solid directly to vapor.” Ice crystals sublime under low humidity in below-freezing conditions. Sublimation fog occurs when ground frost sublimes at sunrise, increasing atmospheric moisture. This can cause a rapid onset of short-lived, shallow, foggy conditions reducing visibility to as low as 1/2 mile. 1.1.2.2.2.7. General Fog Forecasting Guidance: 1.1.2.2.2.7.1. Fog lifts to stratus when the lapse rate approaches dry adiabatic. 1.1.2.2.2.7.2. Marked downslope flow prevents fog formation. 1.1.2.2.2.7.3. The wetter the ground, the higher the probability of fog formation. 1.1.2.2.2.7.4. Atmospheric moisture tends to sublimate on snow, making fog formation, and maintenance less likely. 1.1.2.2.2.7.5. With sufficient radiational cooling (below freezing), fog can dissipate rapidly and form ground frost through the deposition process. 12 AFH15-101 5 NOVEMBER 2019 1.1.2.2.2.7.6. Rapid formation or clearing of clouds can be decisive in fog formation. Rapid clearing at night after precipitation is especially favorable for the formation of radiation fog. 1.1.2.2.2.7.7. The wind speed forecast is important because decreases may lead to the formation of radiation fog. Conversely, increases can prevent fog, dissipate radiation fog, or increase the severity of advection fog. 1.1.2.2.2.7.8. A combination advection-radiation fog is common at stations near warm water surfaces. 1.1.2.2.2.7.9. In areas with high concentrations of atmospheric pollutants, condensation into fog can begin before the relative humidity reaches 100%. 1.1.2.2.2.7.10. The visibility in fog depends on the amount of water vapor available to form droplets and on the size of the droplets formed. At locations with large amounts of combustion products in the air, dense fog can occur with a relatively small water vapor content. 1.1.2.2.2.7.11. After sunrise, the faster the ground temperature rises, the faster fog and stratus clouds dissipate. 1.1.2.2.2.7.12. Solar insolation often lifts radiation fog into thin, multiple layers of stratus clouds. 1.1.2.2.2.7.13. If solar heating persists, and no higher clouds block surface heating, radiation fog usually dissipates. 1.1.2.2.2.7.14. Solar heating may lift advection fog into a single layer of stratus clouds and eventually dissipate the fog if the insolation is sufficiently strong. 1.1.2.2.2.8. In summary, the following characteristics are important to consider when forecasting fog: 1.1.2.2.2.8.1. Synoptic situation, time of year, and station climatology. 1.1.2.2.2.8.2. Thermal (static) stability of the air, amount cooling/moistening expected, wind strength, and dew-point depression. of 1.1.2.2.2.8.3. Trajectory of the air over types of underlying surfaces (i.e., cooler surfaces, bodies of water). 1.1.2.2.2.8.4. Terrain, topography, and land surface characteristics. 1.1.3. Visibility Details. 1.1.3.1. Fog. Fog may develop or intensify when one or more of the following conditions are satisfied: AFH15-101 5 NOVEMBER 2019 13 1.1.3.1.1. Air at the surface is saturated or slightly supersaturated with respect to water in the presence of cloud condensation nuclei. For fog to form, the air near the Earth’s surface must be saturated, or nearly so, meaning that the temperature must equal the dew point (RH = 100%). In reality, because fog is simply a cloud that has formed at the surface, the surface air must be supersaturated for the fog to form. This wouldn’t be possible if the air was perfectly pure with no particulates, but because the air near the surface has many particles of dust, soil, and other minerals, the air can become supersaturated (RH slightly greater than 100%), which allows visible cloud droplets to form. However, mixed-phase (freezing) fogs and ice fogs can develop even if the environment is slightly unsaturated with respect to liquid water (RH less than 100%). Mixed-phase fogs are composed of supercooled water droplets and ice crystals, while ice fog is composed entirely of ice crystals and usually occurs at very cold temperatures (less than -30°C). For the formation of these fogs, RH with respect to water will typically be between 99% and 99.9%, while RH with respect to ice may or may not be greater than 100%. The surface air can become saturated via the following mechanisms: 1.1.3.1.1.1. Temperature cooling to the dew point via cold air advection or radiational cooling. 1.1.3.1.1.2. Dew point increasing to the temperature via moisture advection or evapotranspiration from the earth’s surface. 1.1.3.1.1.3. Precipitation very rapidly raises the dew point and cools the temperature so they eventually equal each other. This process is called “wetbulbing”, whereby precipitation falls through an unsaturated layer of air and the resultant evaporative cooling eventually lowers the temperature to the wet-bulb temperature, while at the same time the dew point increases as moisture is added to the layer. 1.1.3.1.1.4. Introduction of more fine particulates (ash, dust, etc.) into surface air that is already very moist, which allows supersaturation to more easily be achieved. This is akin to cloud seeding. 1.1.3.1.2. A wet ground surface exists, such as a body of water, moist vegetation, or soil recently moistened by precipitation. While not imperative for fog formation, a wet surface such as water or moist vegetation serves as a source of moisture for the planetary boundary layer (PBL). This moisture increases the dew point of the PBL and reduces the amount of cooling that is required for fog formation. A wet surface is especially important for radiation fog formation (discussed in greater detail below), which may not always have an advective source of surface moisture such as from an ocean or lake. 1.1.3.1.3. Surface winds are less than 10 knots. Advection fog is characterized by stronger surface winds than radiation fog, but in general surface winds greater than 10 knots are detrimental to fog formation because they make it difficult for saturation to be achieved by maintaining a well-mixed PBL and bringing in drier ambient air. When fog has already formed, however, the turbulence generated by surface winds less than 10 knots plays a pivotal role in the intensity and maintenance of the fog deck. 14 AFH15-101 5 NOVEMBER 2019 1.1.3.1.4. Dry air with relative humidity less than 50% exists above the surface moist layer. Dry air above the surface is primarily important for radiation fog development and evolution, but is also important for the transition of advection fog to radiation fog. Dry air above the surface to great heights aloft is indicative of clear skies that are conducive to radiative cooling of the earth’s surface. This allows the temperature to cool to the dew point, which also creates a surface temperature inversion. At some point, if the surface temperature cools to the dew point temperature, the air becomes saturated and fog may form. When fog does form, dry air aloft allows the top of the fog deck to continue to cool via outgoing IR radiation, which reinforces the fog deck and often causes it to intensify. The presence of dry air aloft also has an impact on the thickness of the fog deck. For example, because the residual layer (RL) above the surface temperature inversion is usually dry adiabatic, winds within the RL will mix dry air above the inversion into the fog deck. This causes the intensity of the fog deck to fluctuate, but also causes it to deepen as turbulence develops within the fog deck. When the sun rises, dry air aloft permits maximum heating of the fog deck. Solar radiation is able to penetrate to some depth within the fog, which eventually causes it to lift and dissipate when IR cooling is no longer sufficient to maintain the fog. 1.1.3.1.5. Moisture convergence is occurring, especially along coastlines or in areas of complex terrain. Moisture convergence is a measure of the degree to which moist air (water vapor) is converging into a given area. It takes into account converging winds and the advection of moisture by these winds. Moisture convergence occurs along frontal boundaries, complex terrain, and coastlines where the transition from “smooth” water to land with higher friction causes winds to slow down and converge. Moisture convergence is one of the most important aspects of advection fog formation, since it’s common along coastlines when a sea breeze is present and can quickly bring unsaturated air to saturation and keep it saturated as long as onshore flow is maintained. As such, it’s the primary factor in coastal advection fog events, especially if other factors (e.g., subsidence, dry air aloft, etc.) are favorable. 1.1.3.1.6. Precipitation falls into a slightly unsaturated layer, bringing surface relative humidity with respect to water to 100%, or slight supersaturation (relative humidity greater than 100%). Light precipitation (drizzle, mist) often leads to fog formation in the presence of warm fronts, cold fronts, or other convergent boundaries. When light precipitation falls into an unsaturated layer of air, some of it evaporates which cools and moistens the unsaturated air. This simultaneously lowers the temperature and raises the dew point, making precipitation an ideal way to saturate a layer of air. If winds are light and all of the particulates in the air have not been washed to the surface by heavier rain, fog often forms. 1.1.3.1.7. Warm air advection occurs within an existing temperature inversion above a surface moist layer. Warm air advection (WAA) within the temperature inversion above the fog layer is important to fog duration. If WAA is occurring in the inversion, the inversion will be strengthened and will take longer to be mixed out by daytime heating and turbulent mixing. This will allow the foggy air to be trapped under the inversion for a longer period of time. AFH15-101 5 NOVEMBER 2019 15 1.1.3.1.8. The atmosphere is statically or conditionally stable. Stability is one of the most important considerations when forecasting fog. Fog will rarely form in a statically or conditionally unstable PBL, because drier air from aloft is allowed to mix with moist surface air. In other words, the moist surface air is not isolated from the drier air aloft. Most fog events occur under a surface anticyclone (surface high pressure), which effectively holds water vapor at the surface and isolates the surface from the typical well-mixed layer above an inversion. Both temperature inversions and subsidence inversions suppress substantial vertical mixing, and allow shallow fog layers to develop. 1.1.3.1.9. Surface dew points are high. While not imperative for fog formation, a high surface dew point generally means that the surface will not need to cool as much to achieve saturation. In addition, a higher dew point means there is more moisture near the surface, which is important for potential fog thickness. Fogs that form in maritime environments with high dew points will generally be thicker than fogs that form where dew points are lower. 1.1.3.1.10. Upslope flow is occurring. Fog is very common upstream along mountain ranges where moist air impinges on the mountain barrier, is lifted, and condenses to form clouds along the mountain side. Upslope flow is basically moisture convergence along a mountain barrier or other area of complex terrain. If a moist flow pattern persists, then upslope fog can last for several days as moisture is continually lifted along the mountain-side and condenses to form fog. 1.1.3.1.10.1. The opposite is downslope (katabatic) flow, which results in compressional warming of descending air and often leads to fog dissipation during the daylight hours. There are cases, however, where this sinking air along the mountain side spreads out on top of colder air in the valley, reinforcing an existing fog deck by creating a temperature inversion that traps the valley moisture. 1.1.3.1.11. There is snow cover and/or the ground is frozen. Advection fog often occurs during the winter months when warm, moist air moves over snow covered and/or frozen ground. The surface air is cooled from below by the ground surface, which may bring the surface air to saturation, allowing fog to develop. 1.1.3.1.12. There is cloud cover during the day and/or clear skies at night. Cloud cover above a fog layer can cause the layer to remain in place or intensify during the day by absorbing solar radiation that would otherwise penetrate the fog layer and cause it to dissipate. During nighttime hours, however, the same clouds can cause the fog layer to slowly dissipate as the clouds emit down-welling IR radiation into the fog layer, which mitigates some of the cooling of the fog layer from below. 1.1.3.2. Specific Forecasting Guidance – Advection Fog: Advection fog is relatively shallow and accompanied by a surface-based inversion. The depth of this fog increases with increasing wind speed (though at wind speeds above 9 knots greater turbulent mixing usually causes advection fog to lift into a low stratus cloud deck). Other favorable conditions include: 16 AFH15-101 5 NOVEMBER 2019 1.1.3.2.1. Coastal areas where moist air is advected over water cooled by upwelling. During late afternoon, such fog banks may be advected inland by sea breezes or changing synoptic flow. These fogs usually dissipate over warmer land; if they persist through late afternoon, they can advect well inland after evening cooling and last until convection develops the following morning. 1.1.3.2.2. In winter, when warm, moist air flows over colder land. This is commonly seen over the southern or central United States and the coastal areas of Korea and Europe. Because the ground often cools by radiation cooling, fog in these areas is called advection-radiation fog, a combination of radiation and advection fogs. 1.1.3.2.3. Warm, moist air that is cooled to saturation as it moves over cold water forms sea fog. If the initial dew point is less than the coldest water temperature, sea fog formation is unlikely. In poleward-moving air, or in air that has previously traversed a warm ocean current, the dew point is usually higher than the cold water temperature. Sea fog dissipates if a change in wind direction carries the fog over a warmer surface. An increase in the wind speed can temporarily raise a surface fog into a stratus deck. Over very cold water, dense sea fog may persist even with high winds. The movement of sea fog onshore to warmer land leads to rapid dissipation. With heating from below, the fog lifts, forming a stratus deck. With further heating, this stratus layer changes into a stratocumulus cloud layer and eventually changes into convective clouds or dissipates entirely. Cooling after the heat of the day can cause sea fog to roll back in and restrict ceilings and visibility again. 1.1.3.3. Specific Forecasting Guidance – Radiation Fog: Radiation fog occurs in air with a high dew point. This condition ensures radiation cooling lowers the air temperature to the dew point. The first step in making a good radiation fog forecast is to accurately predict the nighttime minimum temperature. Additional factors include the following: 1.1.3.3.1. Air near the ground becomes saturated. When the ground surface is dry in the early evening, the dew- point temperature of the air may drop slightly during the night due to condensation of some water vapor as dew or frost. 1.1.3.3.2. In calm conditions, this type of fog is limited to a shallow layer near the ground; wind speeds of 3-7 knots bring more moist air in contact with the cool surface and cause the fog layer to thicken. A stronger breeze prevents formation of radiation fog due to mixing with drier air aloft. 1.1.3.3.3. Constant or increasing dew points with height in the lowest 200 to 300 feet, so that slight mixing increases the humidity. 1.1.3.3.4. Stable air mass with cloud cover during the day, clear skies at night, light winds, and moist air near the surface. These conditions often occur with a stationary, high-pressure area. 1.1.3.3.5. Relatively long period of radiational cooling, e.g., long nights and short days associated with late fall and winter in humid climates of the middle latitudes. 1.1.3.3.6. In nearly saturated air, light rainfall will trigger the formation of ground fog. 1.1.3.3.7. In valleys, radiation fog formation is enhanced due to cooling from cold air drainage. This cooled air can result in very dense fog. AFH15-101 5 NOVEMBER 2019 17 1.1.3.3.8. In hilly or mountainous areas, an upper-level type of radiation fog— continental high inversion fog—forms in the winter with moist air underlying a subsiding anticyclone. A stratus deck often forms at the base of the subsidence inversion, then lowers. Since the subsiding air above the inversion is relatively clear and dry, air at the top of the cloud deck cools by long-wave radiational cooling, which intensifies the inversion and thickens the stratus layer. A persistent form of continental high inversion fog occurs in valleys affected by maritime polar air. The moist maritime air may become trapped in these valleys beneath a subsiding stagnant high-pressure cell for periods of two weeks or longer. Nocturnal long-wave radiational cooling of the maritime air in the valley causes stratus clouds to form for a few hours the first night after the air becomes trapped. These stratus clouds usually dissipate with surface heating the following day. On each successive night, the stratus cloud deck thickens and lasts longer into the next day. The presence of fallen snow adds moisture and reduces daytime warming, further intensifying the stratus and fog. In the absence of air mass changes, eventually the stratus clouds lower to the ground. The first indicator of formation of persistent continental high inversion fog is the presence of a wellestablished, stagnant high- pressure system at the surface and 700 mb level. In addition, a strong subsidence inversion separates very humid air from a dry air mass aloft over the area of interest. The weakening or movement of the high-pressure system and the approach of a surface front dissipates this type of fog. 1.1.3.3.9. Radiation fog sometimes forms about 100 feet (30 meters) above ground and builds downward. When this happens, surface temperature rises sharply. Similarly, an unexpected rise in surface temperature can indicate impending deterioration of visibility and ceiling due to fog. 1.1.3.3.10. Radiation fog dissipates from the edges toward the center. This area is not a favorable area for cumulus or thunderstorm development. 1.1.3.4. Specific Forecasting Guidance – Frontal Fog. Frontal fog forms from the evaporation of warm precipitation as it falls into drier, colder air in a frontal system. 1.1.3.4.1. Pre-frontal, or warm-frontal, fog (Figure 1.2) is the most common and often occurs over widespread areas ahead of warm fronts. Whenever the dew-point temperature of the overrunning warm airmass exceeds the wet-bulb temperature of the cold airmass it‘s replacing, fog or stratus form. Fog usually dissipates after frontal passage due to increasing temperatures and surface winds. 1.1.3.4.2. Post-frontal, or cold-frontal, fog (Figure 1.3) occurs less frequently than warm-frontal fog. Slow-moving, shallow-sloped cold fronts characterized by vertically decreasing winds through the frontal surface, produce persistent, widespread areas of fog and stratus clouds 150 to 250 miles behind the surface frontal position to at least the intersection of the frontal boundary with the 850 mb level. Strong turbulent mixing behind fast-moving cold fronts, characterized by vertically increasing winds through the frontal surface, often produce stratus clouds but no fog. 18 AFH15-101 5 NOVEMBER 2019 Figure 1.2. Pre-frontal fog associated with a warm front. Figure 1.3. Post-frontal fog associated with a slow-moving cold front. 1.1.3.5. Snow and Blowing Snow. According to the AMS Glossary of Meteorology, blowing snow is snow lifted from the surface of the earth by the wind to a height of 2 meters (6 feet) or more, and blown about in such quantities that horizontal visibility is reduced to less than 7 statute miles. Blowing snow is encoded as BS in surface aviation weather observations and as BLSN as an obstruction to vision in METAR or SPECI observations. Blowing snow can be falling snow or snow that has already accumulated but is picked up by strong winds. Consider the following rules of thumb when forecasting visibility in snow or blowing snow: 1.1.3.5.1. Moderate, dry, and fluffy snowfall with wind speeds exceeding 15 knots usually reduces visibility in blowing snow. 1.1.3.5.2. Snow cover that has previously been subject to wind movement (either blowing or drifting) usually does not produce as severe a visibility restriction as new snow. 1.1.3.5.3. Snow cover that fell when temperatures were near freezing does not blow except in very strong winds. 1.1.3.5.4. The stronger the wind, the lower the visibility in blowing snow. The converse is also true; visibility usually improves with decreasing wind speed. AFH15-101 5 NOVEMBER 2019 19 1.1.3.5.5. Loose snow becomes blowing snow at wind speeds of 10 to 15 knots or greater. Although any blowing snow restricts visibility, the amount of the visibility restriction depends on such factors as terrain, wind speed, snow depth, and composition. 1.1.3.5.6. Blowing snow is a greater hazard to flying operations in polar regions than in mid- latitudes because the colder snow is dry, fine and easily lifted. 1.1.3.5.7. Winds may raise the snow 1000 feet above the ground, lowering visibility. A frequent and sudden increase in surface winds in polar regions may cause the visibility to drop from unlimited to near zero within a few minutes. 1.1.3.5.8. Fresh snow blows or drifts at temperatures of –20°C (–4°F) or less. After 3 or more days of exposure to direct sunlight, snow forms a crust and does not readily drift or blow. The crust, however, is seldom uniform across a snowfield. Terrain undulations, shadows, and vegetation often retard the formation of the crust. 1.1.3.5.9. If additional snow falls onto snowpack that has already crusted, only the new snow blows or drifts. 1.1.3.5.10. Use Table 1.1 as a guide to forecast visibility based on the intensity of snow. When forecasting more than one form of precipitation at a time, or when forecasting fog to occur with the precipitation, consider forecasting a lower visibility than shown in Table 1.1 Table 1.1. Visibility limits based on snowfall intensity. Intensity Visibility Limits (statute miles) Light snow showers Greater than ½ mile Moderate snow showers Greater than ¼ mile but less than ½ mile Heavy snow showers Less than ¼ mile 1.1.3.6. Haze. Research shows that the primary constituent of haze droplets over industrial areas, such as the central and eastern United States and parts of Asia, is sulfuric acid. In these areas, sulfur dioxide released from industry (such as oil refining and steel manufacturing) bonds with oxygen in oxidation reactions enhanced by sunlight and/or liquid water. These reactions result in the formation of sulfate aerosols, include sulfuric acid. Sulfate aerosols are hygroscopic (they absorb water from the environment), even at relative humidities as low as 70%, making them effective condensation nuclei. When the environment is supersaturated, sulfate aerosols grow large enough to be seen as clouds. When the humidity is low, however, they only grow to around a tenth of a micrometer in diameter, and remain suspended in the atmosphere to form haze. Haze usually occurs in the planetary boundary layer (PBL), which extends from the surface up to about 2 km on average, but can extend up to 500 mb in places like Southwest Asia. The top of the PBL is delineated by a temperature inversion and a cessation of vertical mixing, and this is typically the upper boundary of a haze layer. However, elevated layers of haze may also occur, such as in regions downstream from where particles have been lofted above the PBL by cumulus clouds. Elevated haze is also possible on the poleward side of fronts. Sulfate aerosols are chemically stable, and as such, settle very slowly. Therefore, in the absence of a cleansing mechanism such as precipitation, haze can persist for days. This is especially true when atmospheric conditions are stagnant, such as under a persistent area of high 20 AFH15-101 5 NOVEMBER 2019 pressure. Here, limited mixing, plenty of sun, and increasing boundary layer moisture create particularly favorable conditions for the formation of sulfate aerosols. Haze normally restricts visibility to 3 to 6 miles, and occasionally to less than 1 mile. It usually dissipates when the atmosphere becomes thermally unstable or wind speeds increase. This can occur with heating, advection, or turbulent mixing. 1.1.3.7. Dust Storms. Dust storms are a function of wind speed, wind direction and soil moisture content. After generating blowing dust upstream, wind speed becomes important in advection of the dust. Dust may be advected by winds aloft when surface winds are weak or calm. Duration of the advected dust is a function of the depth of the dust and the advecting wind speeds. Synoptic situations, such as cold frontal passages, may change both the wind direction and the probability of dust advecting into your area. 1.1.3.7.1. Dust Source Regions. Forecasting dust generation is more difficult than forecasting the advection of observed dust into the area; there are several key factors to keep in mind, including the following common dust source regions: 1.1.3.7.1.1. Deserts. Dust storms occur in regions with little vegetation and precipitation; these conditions are most prevalent in deserts, where rainfall is scarce. In general, dust is unlikely within 24 to 36 hours of a rainstorm, but rainstorms can lead to increased dust potential beyond 36 hours. Runoff after heavy rain carries soil particles picked up by erosion; once the water evaporates, these particles can be easily lofted. 1.1.3.7.1.2. Agricultural areas. Agricultural land that is fallow, recently tilled, or has a marginal growing climate is a potential source area for dust. The mechanical breaking of soil creates an environment rich in fine-grained soil that is picked up and moved by seasonal winds. This is commonly observed in northern Syria and Iraq. 1.1.3.7.1.3. Coastal areas. Dust plumes can be generated in advance of cold fronts moving across sandy or silty coastal regions. 1.1.3.7.1.4. River flood plains (alluvial plains). The flood plain of the Tigris and Euphrates Rivers in southern Iraq serves as the source region for many dust storms, particularly during shamal events - a northwesterly wind that blows over Iraq and the Persian Gulf states. Shamals are often strong during the day and decrease in strength at night. 1.1.3.7.1.5. Dry lake beds. Water in lakes erodes rocks and forms fine-grained soils. When lakes dry up, these soils inhibit plant growth and blow easily in strong winds. 1.1.3.7.2. Dust Lofting Mechanisms. After an appropriate source, the next key ingredient for dust storm generation is wind from the surface through the depth of the boundary layer strong enough to move and loft dust particles. Table 1.2 shows threshold dust-lofting wind speeds for different desert environments. The first sand and dust particles that move are those from 0.08 to 1 mm in diameter; this occurs with wind speeds of 10 to 25 knots. In general, winds at the surface need to be 15 knots or greater to mobilize dust. Once a dust storm starts, it can maintain the same intensity even when wind speeds slow to below initiation levels, since the bond between dust particles and AFH15-101 5 NOVEMBER 2019 21 the surface is already broken. Lofting of dust also requires turbulence in the boundary layer. Typically, wind shear creates the turbulence that lofts dust up and away from the surface. As a rule of thumb, if the wind at the surface is blowing 15 knots, the wind at 1000 feet) must be about 30 knots to keep the dust particles aloft. An unstable boundary layer, which favors vertical motions, is also necessary for dust lofting. With the lack of vegetation in dust-prone regions, the ground can experience extreme daytime heating, which creates an unstable boundary layer. As the amount of heating increases, the unstable layer deepens. In contrast, stable boundary layers suppress vertical motions and inhibit dust lofting. Diurnal effects also impact dust storm potential. Dry desert air has a wide diurnal temperature difference. Strong radiative cooling leads to rapid heat loss after sunset, resulting in a surface-based inversion, which may inhibit dust lofting. However, the formation of a surface-based inversion has little effect on dust that is already suspended higher in the atmosphere. Furthermore, if winds are sufficiently strong, they will inhibit the formation of an inversion or even remove one that has already formed. Table 1.3 shows favorable parameters for dust storm generation in various scenarios. Table 1.2. Threshold dust-lofting wind speeds for different desert environments. Environment Threshold Wind Speed Fine to medium sand in dune-covered areas 10 to 15 mph (8.7 to 13 knots) Sandy areas with poorly developed desert pavement 20 mph (17.4 knots) Fine material, desert flats 20 to 25 mph (17.4 to 21.7 knots) Alluvial fans and crusted salt flats (dry lake beds) 30 to 35 mph (26.1 to 30.4 knots) Well-developed desert pavement 40 mph (36.8 knots) Table 1.3. Favorable conditions for the generation and advection of dust. Parameter or Condition Favorable When Location with respect to source region Located downstream and in close proximity Agricultural practices Soil left unprotected Previous dry years Plant cover reduced Wind speed 30 knots or greater Wind direction Significant dust source is upstream Cold front Passes through the area Squall line Passes through the area Leeside trough Deepening and increasing winds Thunderstorm Mature storm in local area or generates blowing dust upstream Whirlwind In local area Time of day 1200 to 1900L Surface dew point depression 10 C or greater 1.1.3.7.3. Dust Removal Mechanisms. Lofted dust eventually settles, but may travel half way around the globe before doing so. The three most common dust removal mechanisms are dispersion, advection, and entrainment in precipitation. 22 AFH15-101 5 NOVEMBER 2019 1.1.3.7.3.1. Dispersion. Dust plumes tend to fan out as they move downstream from their source regions; this is caused by dispersion. At the most basic level, the more air that entrains into a dust plume, the more the plume dilutes, spreads out, and disperses. Dispersion is primarily influenced by turbulence, which mixes ambient air with the dust plume. Any increase in turbulence increases the rate at which the plume disperses. There are three kinds of turbulence that act to disperse a plume - mechanical, caused by air flowing over features such as hills or buildings; shear, resulting from differences in wind speed and/or direction; and buoyancy, caused by parcels of air rising during the diurnal heating of the surface. 1.1.3.7.3.2. Advection. Advection moves dust away from its source; winds aloft may carry dust in a direction different from the wind direction at the surface. When predicting where a dust plume will travel, check the vertical wind profile. Dust that leaves the ground going one direction can rise to a level where it travels in an entirely different direction. 1.1.3.7.3.3. Entrainment in precipitation. Most dust particles are hygroscopic, or water-attracting. In fact, dust particles usually form the nucleus of precipitation. Because of this affinity to moisture, precipitation very effectively removes dust from the troposphere. 1.1.3.7.4. Dust storm types. Several types of dust storms, including shamal, frontal, and convective, are described below. The shamal is unique to the Middle East, but frontal and convective dust storms can be experienced in other arid regions, including the Southwestern United States. 1.1.3.7.4.1. Summer shamal. The term “shamal” comes from the Arabic word for “north,” and describes a type of dust storm caused by prevailing north winds over the Arabian Peninsula, Iraq, and Kuwait. Typical sources for the dust lie between the Tigris and Euphrates rivers. The summer shamal, also known as the “Wind of 120 days,” is nearly constant from June through September across Syria, Saudi Arabia, and Iraq. It is caused by convergence between the Southern Arabian Peninsula Monsoon Trough and the subtropical ridge that extends from the Mediterranean Sea into Iraq and the northern Arabian Peninsula. Additionally, cold air advection aloft causes steep lapse rates and increased instability. The end result is that updrafts keep dust particles suspended, sometimes at high altitudes. Due to the dry desert air and high rate of thermal radiation during the day, a nocturnal radiational inversion is established almost every night across Southwest Asia (SWA) as the desert floor cools. The inversion is usually quite shallow, extending up to 1000-2000 feet above ground level. Above the inversion, a nocturnal lowlevel jet (LLJ) stream is often established. When the inversion collapses or is mixed out during the ensuing daylight hours, the strong LLJ winds can reach the surface and are also a source of the summer shamal, as long as they exceed 15 knots. Summer shamal dust storms are typically 3000-8000 feet tall, but can extend up to 15,000-18,000 feet. Visibilities can go from unrestricted to zero within minutes and remain that way for 1-3 hours before slowly starting to increase. Dust storms can last from 1-10 days depending on the wind flow pattern and atmospheric stability profile, but 3 days is average. AFH15-101 5 NOVEMBER 2019 23 1.1.3.7.4.2. Winter shamal. Many winter shamal dust storms are associated with passing cold fronts, which are classified as “frontal dust storms.” A few, however, are caused as very cold air masses funnel south or southeastward from Turkey or Syria into the Tigris/Euphrates River Valley. These cold air masses maintain temperature and moisture continuity as they descend the Arabian Peninsula, creating a relatively sharp temperature gradient between the leading edge of the air mass and the surrounding air. Because the air mass is much colder than the surrounding air, it lifts the warmer and more buoyant ambient air, as well as dust if the ambient vertical temperature profile is conditionally unstable aloft. This physical mechanism, known as a density current, is the same mechanism that lofts dust when a thunderstorm downdraft reaches the surface, causing a haboob. 1.1.3.7.4.3. Frontal dust storms. The main difference between frontal dust storms and shamals is a matter of scale; shamals are localized, mesoscale phenomena, while frontal dust storms are caused by synoptic-scale systems whose winds carry sand particles over large distances. They accompany low pressure systems and associated frontal boundaries. The three major varieties are prefrontal, postfrontal, and shear-line. 1.1.3.7.4.3.1. Prefrontal. Prefrontal dust storms occur across much of SWA as low pressure systems move across the region. The polar front jet stream located behind a cold front and the subtropical jet stream located ahead of a cold front can converge into a single jet streak (or jet max) that translates to the surface in the left front quadrant. Ageostrophic circulations associated with the divergent (left-front and right-rear) quadrants of the jet streak also increase upward vertical velocities, which enhance the probability of dust lofting as well. Prefrontal winds are called the Sharqi in Iraq, the Kaus in Saudi Arabia, the Shlour in Syria and Lebanon, and the Khamsin in Egypt. Easterly to southerly prefrontal winds are favored for prefrontal dust storms in October and November, but such plumes will rarely be long-lived given their prefrontal nature. Wind speeds are 10-20 knots on average, although gusts of 25-30 knots do occur. Prefrontal dust storms can be difficult to detect in METSAT imagery because they are short-lived and often located over similarly shaded terrain. Cloud cover due to pre-cold frontal overrunning often obscures these types of dust storms from METSAT view. 1.1.3.7.4.3.2. Postfrontal. Postfrontal dust storms are associated with a dynamic weather feature (e.g., cold front); cloudiness associated with the cold front may mask the dust signature in METSAT imagery, but in most cases the dust “head” is the best indicator of the leading edge of the cold air. Given their stronger mechanical forcing, postfrontal dust storms usually reach greater heights than those associated with shamal winds. They typically reach 8,00015,000 feet, but postfrontal dust has been observed up to jet stream level (~30,000 feet). Surface winds are 15-30 knots on average, but gusts of 40-50 knots may occur with very strong low pressure systems. Wind speed can be estimated by the thermal contrast across the front. A temperature change of 10°C corresponds to a maximum wind speed of 30 knots, with a 5 knot increase for every additional 5°C. There are two types of postfrontal dust storms: The 24 AFH15-101 5 NOVEMBER 2019 first type lasts 24-36 hours as a cold frontal system migrates across SWA. Dust moves across the Persian Gulf in 12-24 hours, and the cold front can reach the southern Arabian Peninsula in 48-72 hours. The second type lasts 3-5 days and occurs when a weaker cold front (typically in early autumn) stalls out and becomes stationary. Cyclogenesis occurs along the stationary boundary, with frontal waves moving to the east-northeast. 1.1.3.7.4.3.3. Shear line. These dust storms are common in winter along the Arabian Peninsula, Red Sea, and in the equatorial regions of Africa. Trade winds from the east converge with a polar high pressure system, resulting in a narrow band of accelerated flow. This creates enough turbulence to lift dust particles into the atmosphere. Being slightly warmer than the cool air, the dust is sometimes apparent on infrared satellite imagery. A similar situation is observed on the Arabian Sea when a cold front weakens across the IndiaPakistan border. A shear line is produced, which can yield a rope cloud if sufficient moisture is present. Wind speeds are typically 10-25 knots, with gusts to 40 knots. Visibility typically remains 1-3 nautical miles. 1.1.3.7.4.4. Convective dust storms. Convective dust storms are very difficult to forecast, partially because they are typically a small-scale phenomenon (e.g., dust devils may be only tens of meters in width). Microbursts are particularly dangerous for aircraft, given the reduced visibility and unpredictable nature of high winds. 1.1.3.7.4.4.1. Dust devils. Dust devils occur throughout much of the world; they’re created by strong surface heating under clear skies and light winds when the sun warms the air near the ground to temperatures well above those just above the surface layer. Once sufficient heating is achieved, a localized pocket of air quickly rises through the cooler air above it. Hot air rushes in to replace the rising air at the bottom of the developing vortex, intensifying the spinning effect. Once formed, the dust devil is a funnel-like chimney through which hot air moves both upwardly and circularly. The dust devil persists until the supply of warm, unstable air is broken or depleted. Dust devils are typically smaller and less intense than tornadoes, although the strongest dust devils can achieve the intensity of a weak tornado. Dust devil diameter is typically between 10 and 300 feet (3 to 90 meters), with an average height of 500 to 1000 feet (150 to 300 meters). Dust devils typically last only a few minutes, but may persist for up to an hour in optimal conditions. 1.1.3.7.4.4.2. Haboobs. A haboob is an intense dust storm generated by the convective outflow from a collapsing or ongoing thunderstorm, or from any collapsing cumuliform cloud of appreciable vertical extent. Haboobs are most frequent in the deserts of northern Africa, but they also occur in SWA and the southwestern United States. Because low-level air is so dry in desert environments, any precipitation that falls into it will quickly evaporate (or sublimate in the case of ice crystals). This cools the ambient air, making it negatively buoyant with a tendency to sink towards the surface. As long as the wet bulb potential temperature of the sinking air (downdraft) remains cooler than the potential temperature of the air at the surface by at least 4°C, that air will continue to the surface unimpeded and spread out in all directions, but with AFH15-101 5 NOVEMBER 2019 25 a stronger component in the direction of the low-level winds. The air that reaches the surface is colder and therefore denser than the ambient environmental air, so the downdraft air will act like a wedge that rapidly lifts the positively buoyant warm air around it. While this is happening, any dry soil at the cold air/warm air interface has the potential to be lofted to great heights and propagate in the direction of the average cloud-bearing layer winds - this intense dust event is the haboob. The haboob environment is usually characterized by a surge of midlevel moisture, with a dry adiabatic lapse rate below that. Some elevated instability usually exists in the form of a “best lifted index” (the most unstable lifted index) less than 0. Haboobs usually travel an average of 60-90 miles, but can travel greater distances if the convective outflow is strong enough, the low-level winds are strong enough, and/or the vertical temperature profile is dry adiabatic over a large area. Most haboobs reach a height of 5000-8000 feet, but some can ascend to 10,000-14,000 feet. Peak winds in the haboob are approximately 95% greater than the speed of movement. Duration varies, but can reach up to six hours. 1.1.4. Visibility Forecasting Aids and Techniques – Fog. 1.1.4.1. AFW Ensemble Visibility Forecasts. Air Force Weather’s ensemble products (the Global Ensemble Prediction Suite (GEPS) and the Mesoscale Ensemble Prediction Suite (MEPS) use statistical regression analyses of relative humidity, precipitable water, and surface wind speed to produce visibility probability products, available from the AFWWEBS ensembles page at https://weather.af.mil/confluence/display/AFWWEBSTBT/Ensembles+Main+Page. Each ensemble produces probability products for visibilities less than five, three, and one statute miles. 1.1.4.2. Visibility Climatology. The 14th Weather Squadron (https://climate.af.mil) is the Air Force’s climatology center, and operates and maintains a vast library of climatological data. For fog forecasting, their most useful products are wind stratified conditional climatologies (WSCC), surface climograms, and operational climatic data summaries (OCDS-II). 1.1.4.2.1. Wind Stratified Conditional Climatologies (WSCC). Given a set of initial conditions (month, time of day, wind direction, ceiling height, and visibility), WSCCs indicate the percentage likelihood that a particular visibility category will be observed at a future hour (Figure 1.4). The output can be used as a baseline for constructing fog forecasts in areas that are conducive to fog formation. 26 AFH15-101 5 NOVEMBER 2019 Figure 1.4. WSCC example. 1.1.4.2.2. Surface Climograms. A climogram is a two-dimensional view of the likelihood of an event’s occurrence; many weather parameters are available, including fog. Figure 1.5 is a surface climogram for Vandenberg AFB, showing the likelihood of visibility less than 5 statute miles for all hours of the day and all months of the year. Figure 1.5. Visibility Climogram example for Vandenberg AFB. 1.1.4.2.3. Operational Climatic Data Summaries (OCDS-II). The OCDS-II web application enables users to generate plots of fog frequency, stratified by time of day and month of the year, for a user-selected station (Figure 1.6). AFH15-101 5 NOVEMBER 2019 27 Figure 1.6. OCDS-II example. 1.1.4.3. Radiation Fog Point (FP). The radiation fog point is the temperature (in °C) at which radiation fog forms. To find the FP, first find the pressure level of the LCL. From the dew point at the LCL, follow an isohume (line of constant saturation mixing ratio) down to the surface; the temperature value where the isohume intersects the surface is the radiation fog point. For an example of this calculation, refer to Figure 1.7 In the example, the LCL is at 630 mb, and the dew point at the LCL is -40°C. Following the isohume down to the surface from that dew point value, the FP is determined to be -36°C. Figure 1.7. Radiation Fog Point example. 28 AFH15-101 5 NOVEMBER 2019 1.1.4.4. Radiation Fog Threat (FT). The radiation fog threat indicates the potential for radiation fog formation; it’s calculated by subtracting the radiation fog point (see the previous section for details) from the 850 mb wet-bulb potential temperature (WBPT850). If not already known, WBPT850 is found by first finding the pressure level of the 850 mb LCL, and then lowering the parcel moist adiabatically to 1000 mb; the point at which the lowered parcel intersects the 1000 mb level is the WBPT850. Refer to Table 1.4 to determine the likelihood of radiation fog formation. Table 1.4. Fog threat thresholds, indicating likelihood of radiation fog formation. Fog Threat Value Likelihood of radiation fog formation Greater than 3 Low Between 0 and 3 Moderate Less than 0 High Fog Threat Value = WBPT850 – Fog Point 1.1.4.5. Radiation Fog Stability Index (FSI). The radiation fog stability index uses a representative 1200Z sounding to give the likelihood of radiation fog formation, and is defined in Table 1.5 Table 1.5. Fog Stability Index (FSI) thresholds, indicating likelihood of radiation fog formation. FSI Value Likelihood of radiation fog formation Greater than 55 Low Between 31 and 55 Moderate Less than 31 High FSI = 4TSfc - 2(T850 + TdSfc) + W850 TSfc = Surface temperature in °C. T850 = 850 mb temperature in °C. TdSfc = Surface dew point in °C. W850 = 850 mb wind speed in knots. 1.1.4.5.1. The FSI is indicative of radiation fog potential if there’s strong static stability between the surface and 850 mb, as indicated by increasing temperatures with height in the layer. 1.1.4.5.2. FSI is also indicative of radiation fog potential if there’s ample moisture in the layer, indicated by the surface dew point depression (i.e., the difference between the surface temperature and dew point). 1.1.4.5.3. Finally, FSI is also indicative of radiation fog potential if there are slow wind speeds at 850 mb, which aren’t conducive to mixing drier ambient air into the fog layer. 1.1.5. Visibility Forecasting Aids and Techniques – Dust. 1.1.5.1. AFW Ensemble blowing dust forecasts. The GEPS and MEPS produce blowing dust probability products, available on AFW-WEBS, for visibilities less than five, three, and one statute miles. AFH15-101 5 NOVEMBER 2019 29 1.1.5.2. Satellite detection of dust. Satellite detection of dust is difficult, especially at night and in single channel imagery. Enhanced RGB satellite imagery and Aerosol Optical Depth (AOD) have somewhat eased these difficulties. Satellite animations (especially daytime visible imagery) can help identify the location of dust, since the movement of the dust plume makes it stand out against the stationary land surface. Animated infrared imagery can also be used to track dust over land during the day, since the dust contrasts against the hot land surface. Infrared imagery is not as useful at night, since the land cools and the contrast between the dust and land decreases. 1.1.5.2.1. Daytime detection. In general, dust is easier to detect during the day than at night, although there are some differences depending on the time of day. Figure 1.8 depicts visible (A) and infrared (B) Moderate Resolution Imaging Spectroradiometer (MODIS) images, respectively. In the visible image, dust is easily discernible over the Red Sea, but more difficult to observe over the adjoining land; in visible imagery, dust stands out over the dark water background, but blends in with the sandy land surface. The opposite is true in the infrared image; here, dust is clearly depicted over land, but not over the Red Sea. The dust over land is cooler than the underlying surface, and thus detectable on infrared imagery. Over the Red Sea, the thermal contrast is lessened, and the dust disappears. Figure 1.8. Visible (A) and infrared (B) comparison of dust detection capability. 1.1.5.2.2. Sunrise and sunset detection. Dust detection in visible satellite imagery at sunrise and sunset varies from detection during the day. If a satellite is looking in the general direction of the sun and the dust, the forward scattering of dust particles heightens the reflection from dust. Conversely, if the satellite is looking away from the sun, backscattering occurs. Now, less solar energy is being reflected back to the satellite, reducing the ability to detect dust. An example of forward scattering is shown in Figure 1.9; a large dust cloud is seen progressing across the Arabian Peninsula at dawn. This dust plume would be difficult to detect in the middle of the day. 30 AFH15-101 5 NOVEMBER 2019 Figure 1.9. Sunrise dust plume detection, due to forward scattering. 1.1.5.2.3. Aerosol Optical Depth (AOD). AOD is a unit-less measure of the amount of light that airborne particles, such as dust, smoke, haze, and pollution, prevent from passing through a column of the atmosphere. AOD doesn’t provide surface visibility estimates (since the dust detected could reside anywhere in the vertical column), but it can serve as a first-order indicator of potential reduced surface visibilities. Figure 1.10 shows a sample AOD product over Southwest Asia, with warm colors indicating the presence of atmospheric dust. Figure 1.10. AOD product, showing areas of atmospheric dust. AFH15-101 5 NOVEMBER 2019 31 1.1.5.3. Surface Climograms. Climatology can aid in dust forecasting, the surface climogram product can provide a two-dimensional view of the likelihood of reduced visibility due to dust at any hour and month of the year (refer to paragraph 1.1.4.2.2. for details.) 1.1.5.4. Forecasting haboobs from ongoing thunderstorms. First, determine if elevated instability is present (represented by a “best lifted index” less than 0), then look for high mid-level moisture (e.g., high relative humidity between 700 and 500 mb) and steep (dry adiabatic) lapse rates between the surface and approximately 18,000 feet. If all these factors are present, find the strongest wind at any level aloft where the wet bulb potential temperature is less than the surface potential temperature by at least 4°C; this wind may be brought to the surface. 1.1.5.5. Forecasting haboobs from collapsing thunderstorms. Thunderstorm collapse is most likely after sunset when buoyancy diminishes. To determine haboob potential, first find the cloud base height of the thunderstorm. The higher it is (particularly greater than 10,000 feet), the weaker the potential haboob. Rapidly warming cloud tops in infrared satellite imagery are indicative of an imminent collapse. 1.1.5.6. Autumn dust storm forecast process. 1.1.5.6.1. Determine the mission scenario. Determine the time period of interest; you’ll be interested in any dust events that may occur during this time period which may rapidly restrict visibility or negatively impact any missions scheduled or ongoing in your AOR. 1.1.5.6.2. Examine thunderstorm and blowing dust climatology. Refer to AFW-WEBS and the 14th Weather Squadron to determine the climatological likelihood of dust storms at your location. 1.1.5.6.3. Examine a recent dust source region database. Examine the most recent analysis of dust source regions; couple climatology data with historical dust source region information to determine where dust events are most likely to develop. 1.1.5.6.4. Examine a recent soil moisture profile. From AFW-WEBS, examine a 0.25 degree (25 km) resolution 0-10 cm soil moisture profile; locations shaded yellow, light orange, or brown are considered dry. Consult a chart of local soil types determine which soil is most likely to be lofted. Sandy soils drain easier than clay soils and will be more prone to dust lofting, although lofting potential depends strongly on particle size as well. 1.1.5.6.5. Analyze the synoptic environment. At the 200-300 mb level, look for long wave troughs and ridges, as well as jet maxima (or jet streaks). Note the magnitude of jet maxima and areas of divergence in the left front or right rear quadrants. At the 500 mb level, look for shortwave ridges and troughs embedded in the long wave pattern. Also look for areas of positive or negative vorticity advection, and take note of temperatures at 500 mb as indicators of warm or cold air aloft as well as dew points as indicators of dry air aloft. Soundings are also useful when diagnosing mid-level temperatures and moisture. At the 700 mb level, look for shortwave troughs and ridges, and examine dew points. 700 mb temperatures will also indicate any capping inversions that may prevent unimpeded upward vertical motions. At the 850 mb level, look for 32 AFH15-101 5 NOVEMBER 2019 high and low height centers, as well as the presence of a nocturnal low-level jet (LLJ) stream if nightfall is approaching. Also examine 850 mb dew points to determine the amount of low-level dry air present. Analyze the 925 mb level the same the 850 mb level, but key in on wind speeds; 925 mb wind speeds with a dry adiabatic layer often strengthen wind speeds at the surface and enhance dust lofting potential. In fact, the strongest wind anywhere in a dry adiabatic layer could come to the surface and contribute to downdraft strength. Finally, at the surface, analyze areas of high and low pressure and associated fronts, as well as areas of low-level convergence. Surface winds of at least 15 knots are usually needed to loft dust. 1.1.5.6.6. Analyze observed or model forecast soundings. Look for observed or forecast surface winds greater than or equal to 15 knots. In Iraq, north or northwest low-level winds are favorable for dust lofting because they are parallel to several dust source regions. Dry low-level air (from the surface up to somewhere in the 700-400 mb layer) is often conducive to momentum transfer from aloft that can enhance surface winds. Look for the presence (or lack of) a temperature inversion; dust will generally start to settle at the surface after dark once the nocturnal radiational inversion develops. An inversion also makes it difficult for dust to be lofted in the first place due to a shallow layer of stability that is not conducive to upward motion. The presence (or lack of) a dry adiabatic (or super-adiabatic) layer from the surface to some height aloft is a key indicator of dust lofting potential; the height of a dust plume can be approximated by the depth of the dry adiabatic layer. When the temperature becomes less than dry adiabatic, it’s difficult for dust to continue being lofted to greater heights because the potential temperature starts to increase rather than remain constant. It’s often important for winds within the dry adiabatic layer to be in-phase (i.e., coming from the same direction); this increases the probability that winds from aloft will be brought to the surface via momentum transfer. The strongest wind within a dry adiabatic layer may be able to come to the surface and contribute to non-convective or convective wind gusts, which will enhance dust lofting potential. 1.1.5.6.7. Determine the type of dust storm (shamal, frontal, or convective) that’s likely to occur, and make and verify your forecast. Couple the knowledge gained from previous steps with knowledge of the different varieties of dust events that may occur; the general characteristics of the major dust storm types were detailed earlier in the chapter. Forecast if and when dust will be lofted, and when it will settle; the height of a dust plume can be approximated by the maximum height of the dry adiabatic layer in the environment in which the plume occurs, and dust settles at a rate of 1000 feet per hour once surface winds drop below 15 knots or an outflow boundary has passed. Verify your forecast using surface observations and METSAT data. 1.1.5.7. Summer shamal forecast process. 1.1.5.7.1. Determine the mission scenario. Determine the time period of interest; you’ll be interested in any dust events that may occur during this time period which may rapidly restrict visibility or negatively impact any missions scheduled or ongoing in your AOR. AFH15-101 5 NOVEMBER 2019 33 1.1.5.7.2. Examine thunderstorm and blowing dust climatology. Refer to AFW-WEBS and the 14th Weather Squadron to determine the climatological likelihood of dust storms at your location. 1.1.5.7.3. Conduct sounding analysis. Examine a morning sounding within or near dust source regions; note the strength of any inversions and determine if they will break due to turbulent mixing and/or daytime heating. Look for winds that are nearly unidirectional or in-phase through the portion of the troposphere where the lapse rate is dry adiabatic. Add 5-10 knots to low-level wind forecasts if winds are in-phase up to 700 mb (with a dry adiabatic lapse rate), and add 10-15 knots if winds are in-phase up to 500 mb (with a dry adiabatic lapse rate). Note that the strongest wind within the dry adiabatic layer may be able to be brought to the surface in the form of non-convective surface wind gusts. The height of an elevated dust layer can be approximated by determining where the lapse rate becomes less than dry adiabatic. 1.1.5.7.4. Conduct surface analysis. Look for surface wind speeds greater than or equal to 15 knots over relevant source regions; remember to add 5-15 knots to your surface wind speed forecast if winds are in-phase up to 700 mb or 500 mb and the layer is dry adiabatic. Also remember that the strongest wind speed within the dry adiabatic layer can be brought to the surface. Look for wind speeds at 1000 feet AGL (925 mb is a good approximation) greater than or equal to 26 knots over southern Iraq and greater than or equal to 30 knots over Bahrain. 1.1.5.7.5. Conduct METSAT imagery analysis. Use single channel or, preferably, multi-spectral imagery to determine the extent and location of any existing dust plumes. 1.1.5.7.6. Make and verify your dust storm forecast. Forecast the onset, duration, and persistence of any dust events in your area of responsibility during the proposed time period of mission execution (often in the 6-24 hour time period). Ensure you note any local rules of thumb about dust advection, as well as any geographic features such as small dust source regions, terrain, vegetation, and water sources. Note where precipitation has fallen in the past 48 hours; it’ll be much more difficult for dust to be lofted in these areas than in areas with dry soil characteristics. 1.1.5.7.6.1. Verify your forecast by examining METSAT imagery during the mission window and from PIREPS, if available. 1.1.5.8. Haboob forecast process. 1.1.5.8.1. Determine the mission scenario. Determine the time period of interest; you’ll be interested in any dust events that may occur during this time period which may rapidly restrict visibility or negatively impact any missions scheduled or ongoing in your AOR. 1.1.5.8.2. Examine thunderstorm and blowing dust climatology. Refer to AFW-WEBS and the 14th Weather Squadron to determine the climatological likelihood of dust storms at your location. 34 AFH15-101 5 NOVEMBER 2019 1.1.5.8.3. Examine a recent dust source region database. Examine the most recent analysis of dust source regions; couple climatology data with historical dust source region information to determine where dust events are most likely to develop. Areas where thunderstorms are most likely and that are located within or downstream of dust source regions have the highest climatological haboob potential. 1.1.5.8.4. Conduct instability analysis. Using either observed soundings or model data, determine if any elevated instability will be present for updraft development. If using observed soundings without model data, lift the Most Unstable (MU) parcel in the lowest 300 mb of the atmosphere dry adiabatically to its Lifting Condensation Level (LCL). Above the LCL, lift the parcel moist adiabatically to the top of the sounding. Calculated the Lifted Index (LI) by subtracting the temperature of the environment at 500 mb from the temperature of the parcel at 500 mb. A negative LI (less than -2) indicates that convection is possible later in the day. If using model data, examine forecast charts of the Best Lifted Index. 1.1.5.8.5. Conduct mid-level moisture analysis. Diagnose negatively buoyant downdraft potential by examining the contribution due to high water droplet and/or ice crystal concentration, as well as resultant precipitation. Using either observed soundings or model data, determine the maximum relative humidity (RH) in the 700 mb to 400 mb layer. Precipitation aloft is possible where RH is greater than 90% in the 700-400 mb layer. Water droplets falling into a deep, dry planetary boundary layer (PBL) will evaporate and cool the PBL; this is often seen as virga falling from highbased cumulus clouds. At temperatures below -5°C in the 700-400 mb moist layer, a percentage of water vapor will form as ice crystals, further cooling the PBL as they fall into it. Therefore, the higher the RH (greater than 90%) at temperatures below -5°C in the 700-400 mb layer, and the colder the temperatures, the stronger the potential downdraft. If using model data, examine 700 mb and 500 mb forecast charts that plot RH, or examine forecast sounding data in the 700-400 mb layer. 1.1.5.8.6. Examine surface to 500 mb lapse rates. Diagnose negatively buoyant downdraft potential by examining the contribution to downdraft strength due to cold temperatures aloft. Using observed soundings, determine the rate of environmental temperature decrease from the surface to 500 mb (approximately 18,000 feet, or 5.5 km). Subtract the environmental temperature at 500 mb from the surface temperature, and divide this temperature difference by 5.5 km to get an average lapse rate in the surface to 500 mb layer (units: °C/km). The steeper the environmental lapse rate (greater than 5°C/km, with 9.8°C/km optimal), the more negatively buoyant and stronger the potential downdraft. Dust is most easily lofted where the lapse rate within the PBL is dry adiabatic. The maximum height a dust plume can achieve is approximated by the height at which the environmental lapse rate ceases to be dry adiabatic (i.e., less than 9.8°C/km). AFH15-101 5 NOVEMBER 2019 35 1.1.5.8.7. Find the strongest wind aloft that could be brought to the surface. If wetbulb potential temperatures aloft are less than the potential temperature at the surface by at least 4°C, the strongest wind at any of those levels may be able to be brought to the surface, contributing to downdraft strength. Convective downdrafts that result in haboobs will remain negatively buoyant as long as this temperature differential is maintained; the 4°C difference is required because downdrafts may punch through a shallow layer that is somewhat stable to downward motion (i.e., where the wet bulb potential temperature is slightly greater than the surface potential temperature). 1.1.5.8.8. Determine if the mission environment is conducive to haboobs. Assess haboob potential based on the analyses conducted in the previous steps, and incorporate dust model data where applicable. With current technology, it’s not possible to definitively forecast when and where individual haboobs will develop. But using this method, you can determine if the haboob threat in the mission area is minimal, slight, moderate, or high. 1.1.6. Visibility Forecasting Aids and Techniques. 1.1.6.1. Haze. The Air Force’s GEPS and MEPS ensembles produce a Pollution Trapping Index (PTI), which determines the potential for reduced visibility due to haze – the PTI probability products are available from the ensembles main page on AFW-WEBS. Favorable values for pollution trapping are achieved when lapse rates are stable (negative) and winds are light. Under these conditions, pollutants such as sulfuric acid are likely to linger and thicken, producing haze. PTI is calculated using the lapse rate and wind, based on the following algorithm and interpreted according to Table 1.6. Table 1.6. Interpretation of the GEPS and MEPS Pollution Trapping Index (PTI) index. PTI Value Pollution Trapping Potential Greater than -40 Minimal Greater than -30 Slight Greater than -20 Moderate Greater than -10 High PTI = -1.0 X (LAPSE + WIND2) LAPSE = Surface to 700 mb lapse rate in Kelvin/km (positive = unstable) WIND = 10 meter surface wind (m/s) 1.1.6.2. Snow. The GEPS and MEPS produce two blowing snow products, available on AFW-WEBS: joint probability of wind gusts greater than or equal to 25 knots and six-hour snow accumulation greater than 0.1 inch, and joint probability of wind gusts greater than or equal to 35 knots and six-hour snow accumulation greater than one inch. Although these products don’t directly indicate the probability of reduced visibility due to blowing snow, they do indicate the probability of strong winds and accumulating snow occurring concurrently, leading to reduced visibility. 1.1.7. Final thoughts on visibility forecasting. Experience plays an important role in visibility forecasting – consider the following: 36 AFH15-101 5 NOVEMBER 2019 1.1.7.1. Actual Prevailing Visibility. A drop in visibility (i.e., from 25 miles to 15 miles) could indicate a significant increase in low-level moisture that could go unnoticed in a 7+ mile report. 1.1.7.2. Sector Visibility. If sector visibility is significantly different from prevailing, it could mean something significant is occurring. For example, the lowering of sector visibility could mean a fog bank is forming or that dust is rising due to an increase in winds from a thunderstorm. 1.1.7.3. Obstructions to Visibility. Reports should include what is obstructing vision (i.e., fog, smoke, haze, etc.) as well as an estimated layer height top and/or base. For example, “visibility 10 miles in haze, top of haze layer approximately 1500 feet,” includes haze as being the obstruction to vision and identifies the layer of haze. 1.1.7.4. Tops and Bases of Haze Layers. These are important because they may indicate the bases of inversions. Tops and bases of haze layers are usually difficult to estimate, but a definite top and/or base is sometimes detectable when looking towards the horizon. Determine the height by noting the orientation to higher terrain, trees, or buildings, if available. Pilot reports of haze tops and/or bases are also useful. 1.2. Precipitation. For precipitation to occur, two basic ingredients are necessary: moisture and a mechanism for lifting the air sufficiently to promote condensation. Lifting mechanisms include convection, orographic lifting, and frontal lifting. There are many techniques and methods available for forecasting precipitation. 1.2.1. Precipitation General Guidance. 1.2.1.1. Extrapolation. Extrapolation works best in short-period forecasting, especially when precipitation is occurring upstream of the station. First, outline areas of continuous, intermittent and showery precipitation on an hourly or 3-hourly surface product. Use radar and satellite data to refine the surface chart depiction. Use different types of lines, shading, or symbols to distinguish the various types of precipitation. Next, compare the present area to several hourly (or 3-hourly) past positions. If the past motion is reasonably continuous, make extrapolations for several hours. (Note: Consider local effects that may block or slow the movement of the extrapolated area.) 1.2.1.2. Cloud Top Temperatures. The thickness of the cloud layer aloft and the temperatures in the upper levels of clouds are usually closely related to the type and intensity of precipitation observed at the surface, particularly in the mid- latitudes. Monitor satellite imagery to determine if cooling cloud tops are occurring (indicating upward vertical motion). In general, colder cloud tops correspond to a greater chance of precipitation. 1.2.1.3. Dew Point Depression. An upper level dew point depression less than or equal to 2°C is a good predictor of both overcast skies and precipitation. Dew point spreads less than or equal to 2°C on the 850 and 700 mb forecast products are a good indication of potential precipitation, assuming there is also upward vertical motion. AFH15-101 5 NOVEMBER 2019 37 1.2.1.4. Overrunning. Overrunning precipitation occurs in association with active warm fronts, stationary fronts and, to a lesser degree, with slow-moving cold fronts (Figure 1.11). Stratus is a by-product and generally results from the evaporation of relatively warm precipitation into cooler air. Use the 925 mb, 850 mb and 700 mb products to determine whether sufficient moisture and vertical motion are present to produce overrunning precipitation: 1.2.1.4.1. The 925 mb and 850 mb products reveal whether the wind flow is favorable for the advection of this moisture into the area. 1.2.1.4.2. The 700 mb product reveals if the thermal structure is adequate to produce overrunning precipitation. In general, overrunning requires warm-air advection and cyclonic curvature at 700 mb to produce significant precipitation. Therefore, the outer limits of overrunning precipitation are usually the 700 mb ridge line in advance of the system (beginning of precipitation) and behind the system where the wind changes from veering with height (warm- air advection) to backing with height (cold-air advection and the ending of precipitation). Figure 1.11. Overrunning associated with a typical cyclone. 1.2.1.5. Drizzle Formation. The basic requirements for significant drizzle are: 1.2.1.5.1. A cloud layer or fog at least 2000 feet deep. 1.2.1.5.2. Cloud layer or fog must persist several hours to allow droplets time to form. 1.2.1.5.3. Sufficient upward vertical motion to maintain the cloud layer or fog. 1.2.1.5.4. A source of moisture to maintain the cloud or fog. Light drizzle can fall from radiation and sea fog without the help of upward vertical motions. 1.2.1.5.5. Identify areas of potential upward vertical motion by drawing streamlines on surface charts to locate and track local axes of confluence. Make a reasonable estimate of whether surface confluence is stronger or weaker than usual. Drizzle onset is faster and more likely with stronger confluence. 38 AFH15-101 5 NOVEMBER 2019 1.2.1.5.6. Upslope flow and/or sea breeze confluence can produce the weak upward vertical motion required for drizzle formation, without surface observations indicating local confluence. Similarly, persistent large-scale southerly flow naturally converges as it moves northward, and can also provide the required low-level upward motion. Finally, the lift associated with a nearby front can supply the upward motion to generate large areas of fog and stratus. In these instances, it’s possible to extrapolate the onset of drizzle at your location from upstream stations. If surface temperatures are less than or equal to 0°C (32°F), forecast freezing drizzle. 1.2.2. Determining Precipitation Type. 1.2.2.1. Thickness. Thickness (the vertical distance between two constant-pressure surfaces) is the most common predictor for precipitation type. Thickness is a function of temperature: the warmer the air, the thicker the layer. If the thickness of the layer is known, then the layer’s mean temperature can be determined. The most used 1000-500 mb thickness value for forecasting precipitation type is the 540 (5400 meter) threshold. Another predictor is the 0°C 850 mb isotherm. A third predictor is the 850-700 mb, 1550meter thickness line. Studies have shown that snow is rare when the 850-700 mb thickness is greater than 1550 meters, or the 1000-500-mb thickness is greater than 5440 meters. 1.2.2.1.1. Analyzing and Extrapolating Thickness Patterns – Method 1. This method requires that both the low- and mid-level thickness be calculated and plotted, but the precipitation analysis is rapid and straightforward. Using forecast charts, plot the following parameters: 1.2.2.1.1.1. Mid-level thickness (700 mb height minus the 850 mb height). 1.2.2.1.1.2. Low-level thickness (850 mb height minus the 1000 mb height). 1.2.2.1.1.3. 700 mb height contours. 1.2.2.1.1.4. 700 mb dew points. 1.2.2.1.1.5. 850 mb dew points. 1.2.2.1.1.6. The surface 0°C (32°F) isotherm. 1.2.2.1.1.7. The 850 mb 0°C (32°F) isotherm. 1.2.2.1.1.8. Analyze the mid-level thickness for the 1520- and 1540-meter contours and analyze for these dew points: –5°C (850 mb) and –10°C (700 mb). Forecast two or more inches of snow to occur in the area within these lines where precipitation is expected (see Figure 1.12). Analyze the mid-level thickness for the 1555-meter line. Forecast freezing precipitation to occur in the area between this line and the 1540-meter line and within the above dew-point lines, provided the surface temperature is below freezing. Find any areas of appropriate thickness but lacking sufficient moisture at either 850 mb or 700 mb. Be alert for any changes in the moisture pattern by advection or vertical motion. Expect only liquid precipitation on the warm side of the 850 mb 0°C (32°F) isotherm. AFH15-101 5 NOVEMBER 2019 39 Figure 1.12. Method 1 – the plot shows where to expect different precipitation types. 1.2.2.1.2. Analyzing and Extrapolating Thickness Patterns – Method 2. This method requires extrapolation of analyzed thickness contours to their appropriate position at the valid time of the precipitation forecast. The following are general rules for this method of extrapolating thickness patterns: 1.2.2.1.2.1. Low-Level Thickness. Choose several 1000-to-850 mb thickness lines that give a good estimate of the thickness pattern; e.g., the 1300-, 1340-, or 1380meter lines. Move each line in the direction of the wind at 3000 feet with 100 percent of that wind speed. The thickness ridge moves at the speed of the associated short wave. In a strongly baroclinic situation, it moves slightly to the left of the 500 mb flow at 50 percent of the wind speed. Since thickness patterns merely depict the large-scale mass distribution, take care to adjust for rapid changes at 500 mb. Compare the thickness analysis with the surface analysis to ensure a reasonable forecast product. 1.2.2.1.2.2. Mid-Level Thickness. Move the 1520- to 1540-meter band at 100 percent of the 8000-foot wind field. Consider continuity, the latest surface analysis, and other charts when developing a new thickness forecast chart. 1.2.2.1.2.3. Snowfall begins with the approach of a low-level thickness ridge after the passage of the 700 mb ridge line or the line of no 12-hour temperature change (the zero isallotherm) and with the approach of the low-level thickness ridge. Snowfall usually ends after the passage of the low-level thickness ridge and the 700 mb trough. Snowfall is heaviest 1 to 2 hours beforehand, and ends after the passage of the low-level thickness ridge and the 700 mb trough. 1.2.2.2. Freezing Precipitation Indicators. 40 AFH15-101 5 NOVEMBER 2019 1.2.2.2.1. Height of the Freezing Level. Forecasters often use the freezing level to determine the type of precipitation (see Table 1.7). The forecast is based on the assumption that the freezing level must be lower than 1200 feet above the surface for most of the precipitation reaching the ground to be snow. However, forecasters must understand the complex thermodynamic changes occurring in the low levels to correctly forecast winter precipitation situations. For example, the freezing level often lowers 500 to 1000 feet during first 1.5 hours after precipitation begins, due to evaporation or sublimation. When saturation occurs, these processes cease and freezing levels rise to their original heights within 3 hours. With strong warm air advection, the freezing level rises as much as a few thousand feet in a 6- to 8-hour period. The following methods account for both single or multiple freezing levels to forecast the type of precipitation expected at the surface. Each method considers the change of state of precipitation from liquid-to-solid or solid-to-liquid as it falls through the atmosphere. Table 1.7. Probability of snowfall as a function of freezing level height. Height of freezing level above ground 12 mb 25 mb 35 mb 45 mb 61 mb Probability precipitation will fall as snow 90% 70% 50% 30% 10% 1.2.2.2.1.1. Single Freezing Level. If the freezing level equals or exceeds 1200 feet above ground level (AGL), forecast liquid precipitation. If the freezing level is less than or equal to 600 feet AGL, forecast solid precipitation. If the freezing level is between 600 and 1200 feet AGL, forecast mixed precipitation. 1.2.2.2.1.2. Multiple Freezing Levels. When there are multiple freezing levels, warm layers exist where the temperature is above freezing. The thickness of the warm and cold layers affects the precipitation type at the surface. If the warm layer is greater than 1200 feet thick and the cold layer closest to the surface is less than or equal to 1500 feet thick, forecast freezing rain. Conversely, if the warm layer is greater than 1200 feet thick and the cold layer closest to the surface is greater than 1500 feet thick, forecast ice pellets. Finally, if the warm layer is between 600 and 1200 feet thick, forecast ice pellets regardless of the height of the lower freezing level. 1.2.2.2.2. Forecasting Snow vs. Freezing Drizzle. This technique is based on the precipitation nucleation process, and applies to the continental United States, Europe, and the Pacific theater. The technique below assumes the atmosphere is below freezing through its entire depth, and the water droplets remain supercooled until surface contact. 1.2.2.2.2.1. Does a lower-level moist layer (below 700 mb) extend upward to where temperatures are –15°C? If not, then freezing drizzle is possible. 1.2.2.2.2.2. Is a mid-level dry layer (800 to 500 mb) present or forecast? If yes, freezing drizzle or a mixture of snow and freezing drizzle is possible. AFH15-101 5 NOVEMBER 2019 41 1.2.2.2.2.3. Is the mid-level dry layer (dew point depression greater than or equal to 10°C) deeper than 5000 feet? If yes, precipitation may change to freezing drizzle, or a prolonged period of mixed snow and freezing drizzle is possible. 1.2.2.2.2.4. Is mid-level moisture increasing? If freezing drizzle is occurring and mid-level moisture is increasing, precipitation may change to all snow. 1.2.2.2.2.5. Is elevated convection occurring or forecast to occur? If yes, the midlevel dry layer may be eroded, causing snow instead of freezing drizzle. 1.2.3. Freezing Precipitation. Detailed descriptions of the three primary types of freezing precipitation are presented below. 1.2.3.1. Freezing rain. Freezing rain is rain that falls in liquid form but freezes upon contact with sub-freezing surfaces. In the majority of cases, rain must be supercooled (i.e., below freezing) before striking the surface. The two factors that exert the greatest influence on freezing rain potential are the depth and average temperature of the melting layer. Figure 1.13 is a graphical summary of this relationship, and an example of how these factors are primary influences on ice pellet (or sleet) development as well. The figure shows that precipitation type will be ice pellets regardless of melting layer temperature when its depth is less than about 400 meters. Similarly, when the melting layer is greater than 4000 meters deep, precipitation type will almost always be freezing rain, regardless of the layer’s average temperature. Figure 1.14 is an example of a sounding environment in which freezing rain would be observed. Freezing rain will rarely be observed at surface temperatures colder than -10°C (14°F), since too many ice crystals are present for liquid (supercooled) water to remain the predominant hydrometeor type. In fact, most freezing rain events occur with surface temperatures between 0°C (32°F) and -5°C (23°F). Freezing rain usually occurs when the air is very moist (relative humidity [RH] with respect to water greater than 90%) from the surface through 700 mb and into the dendritic layer. This ensures that snowflakes will develop in the dendritic layer. These snowflakes may melt as they fall through the melting layer, becoming supercooled raindrops. The depth of the refreezing layer is also an important consideration, but less so than the above factors. The deeper the refreezing layer, the more likely that a liquid raindrop will become frozen again and reach the surface as sleet or snow rather than freezing rain. The RH within the melting layer must be considered as well. The lower the RH, the greater the distance required for complete melting of snowflakes aloft given the evaporative cooling component. A snowflake falling through a melting layer with 90% RH may take an additional 100 meters to melt than if that same layer was saturated (RH = 100%). Finally, the size and type of snowflakes falling into a melting layer play a role in the melting layer depth and average temperature required for complete snowflake melting. Larger snowflakes (those composed of stellar dendrites or sector plates, for example) require a larger melting layer depth and higher average temperature to melt completely before falling into the refreezing layer. Figure 1.15 is a nomogram of precipitation type based on the 1000 mb wet bulb temperature (°C) and the 850 mb wet bulb temperature (°C), and Figure 1.16 utilizes 1000 to 850 mb thickness (m) and 850 to 700 mb thickness (m) to estimate precipitation type. 42 AFH15-101 5 NOVEMBER 2019 Figure 1.13. Melting layer depth vs. mean layer temperature for freezing rain potential. AFH15-101 5 NOVEMBER 2019 43 Figure 1.14. Typical freezing rain sounding. Note the depth and magnitude of the melting layer, as well as the moisture into the dendritic layer. Figure 1.15. Precipitation type nomogram, based on 1000 mb and 850 mb wet bulb temperatures. 44 AFH15-101 5 NOVEMBER 2019 Figure 1.16. Precipitation type nomogram, based on 1000-850 mb and 850-700 mb thicknesses. 1.2.3.2. Freezing drizzle. Freezing drizzle is precipitation in the form of small, liquid water droplets that freeze upon contact with a surface that is colder than 0°C. Like freezing rain drops, freezing drizzle droplets must be supercooled (colder than 0°C) before reaching the surface. Figure 1.17 shows examples of sounding environments in which freezing drizzle would be expected. Precipitation type will rarely be freezing drizzle if in-cloud and surface temperatures are below -10°C. Ice particles become much more common at temperatures colder than -10°C and tend to deplete small supercooled water droplets necessary for drizzle formation, so frozen precipitation such as light snow or even sleet become more likely when temperatures fall below -10°C. While freezing rain and sleet have deep moisture profiles with high RH extending into the -10°C to -20°C dendritic layer, the moisture associated with freezing drizzle is almost always found below 700 mb, with very dry air aloft. As is apparent in Figure 1.17, the entire moisture profile need not be below freezing for freezing drizzle to be observed at the surface. The top of the cloud can be above freezing as long as the surface layer is below freezing so the drizzle droplets that have developed become supercooled before they reach the surface. Freezing drizzle is less likely in desert or continental environments with copious amounts of dust that can serve as cloud condensation nuclei (CCN). A lack of CCN allows the distribution of small water droplets to be quite broad, increasing the efficiency of the collision-coalescence process by which water droplets develop, and making it easier for drizzle droplets to form. Therefore, freezing drizzle becomes much less likely when cloudy air is filled with CCN. Freezing drizzle almost always falls from stratiform clouds, and the stratiform cloud depth (or thickness) must be at least 200 meters for the microphysical processes to function efficiently. The deeper the stratiform cloud deck, the longer the freezing drizzle is likely to AFH15-101 5 NOVEMBER 2019 45 last, although the cloud deck can’t be so deep as to be seeded by ice crystals that deplete supercooled water droplets. As with freezing rain and sleet, there must be a source of lift (i.e., vertical motion) for freezing drizzle to develop, but only weak vertical motion (5-10 cm/s) is required for drizzle as opposed to the stronger vertical motions that would be required for freezing rain or sleet. In the freezing drizzle sounding to the left in Figure 1.17, the moist layer extends up to 750 mb and is warmer than -6° C. In the freezing drizzle sounding on the right in Figure 1.17, an above-freezing cloud layer extends from 925 mb to 750 mb, but the surface-to-925 mb layer is below freezing, allowing some abovefreezing droplets to become supercooled before impacting the sub-freezing surface. Figure 1.17. Freezing drizzle soundings. 1.2.3.3. Sleet. Sleet (a.k.a. ice pellets) is precipitation in the form of transparent or translucent pieces of ice less than 5 millimeters (mm) in diameter. The two factors that exert the greatest influence on sleet potential are the depth and average temperature of the melting layer (refer back to Figure 1.13 for a graphical summary of this relationship). Precipitation type will be sleet regardless of the melting layer temperature when its depth is less than about 400 meters. In this situation, ice particles that fall into the melting layer don’t have enough time to melt because the layer is too shallow, or the mean temperature of the layer is too cold. Note in Figure 1.13 that the distribution of mean melting layer temperatures and melting layer depth for sleet is much less than for freezing rain, an illustration of the rarity of sleet vs. freezing rain and snow. Sleet is often the “transition” precipitation between freezing rain and snow as the magnitude and depth of the melting layer diminishes or changes with temperature advection and microphysical processes. Figure 1.18 is an idealized sounding environment in which sleet would be observed. The melting layer is relatively deep in this case, but barely above freezing. Also note that moisture extends into the dendritic layer. Unlike freezing drizzle and freezing rain, sleet can occur with surface temperatures well below freezing in an environment with a very strong temperature inversion. As with freezing rain, sleet requires saturated or near saturated conditions (RH greater than 90%) from the surface past 700 mb and into the dendritic layer. If the dendritic layer is not supersaturated with respect to ice (RH greater than 100%), ice crystals will not grow and none will be present to fall into the melting layer. 46 AFH15-101 5 NOVEMBER 2019 Figure 1.18. Idealized sleet sounding. 1.2.3.4. Factors affecting freezing precipitation. The following factors should also be considered when determining the potential for freezing precipitation: 1.2.3.4.1. Cloud Condensation Nuclei (CCN). CCN are particulates in the atmosphere such as dust, minerals, and even ice crystals around which cloud droplets develop (or nucleate). The air in the troposphere is never 100% pure with no CCN. If no CCN existed, it would be very difficult for clouds and precipitation to develop at all. The type, size, and distribution of CCN in a column of air are the most important considerations for precipitation type and intensity. While high concentrations and relatively large CCN are favorable for heavier freezing precipitation such as sleet and/or freezing rain, a lower concentration of CCN is necessary for freezing drizzle formation because the droplet size distribution is broadened, allowing efficient collision and coalescence (or “sticking together”) to occur. Therefore, in regions of high CCN concentrations such as near deserts or other continental areas where dust, soil, and mineral particles are abundant, freezing rain and sleet are more likely than freezing drizzle on most occasions. This is why freezing drizzle often occurs when a shallow, cold air mass replaces a warmer one, pushing away the low-level CCN. 1.2.3.4.2. Cloud depth. Cloud depth is the thickness of a cloud from its base to its top. For cumuliform clouds, the cloud base is assumed to occur at the height of the lifting condensation level (LCL), derived from a surface-based parcel. The parcel is then lifted moist adiabatically from the LCL to the point where the parcel temperature equals or becomes colder than the actual temperature of the atmosphere. This is called the equilibrium level (EL). The EL is assumed to be the cumuliform cloud top, although strong updrafts may allow a portion of the cumuliform cloud to “overshoot” the EL. AFH15-101 5 NOVEMBER 2019 47 The depth of the cloud is found by subtracting the height of the cloud top from the LCL height above the ground. Because stratiform clouds, which often produce freezing precipitation, rarely form due to vertical motions from positive buoyancy, their depth must be estimated differently. Stratiform cloud depth can be directly estimated from the moisture characteristics of the sounding. This is done by using varying values of the dew point depression (DD), which is the difference between the temperature and the dew point, from the surface to some point aloft where the DD thresholds are no longer satisfied. A smaller DD indicates a greater potential that clouds exist at a given level. In general, the larger the cloud depth, the heavier any freezing precipitation that develops will be. Freezing rain, sleet, and snow require much larger cloud depths than freezing drizzle. However, the duration of freezing drizzle correlates closely with the depth of the cloud from which it is falling. Freezing drizzle duration increases as cloud depth increases past 200 m, but this depends on other factors as well, such as temperature and CCN concentration. Freezing rain, sleet, and snow soundings often indicate saturated or near saturated conditions from the surface up to 500 mb (through the -10°C to -20°C dendritic layer), while freezing drizzle soundings may only be saturated from the surface up to 850 mb. 1.2.3.4.3. Ground temperature. The ground temperature can warm or cool the temperature of the lowest few centimeters of the atmosphere via a process called conduction, but other than that, has little influence on the character of falling precipitation. It does, however, exert an enormous influence on the character of precipitation once it hits the ground. Precipitation may fall as rain, but be characterized as freezing rain if the ground temperature is 0°C or less. On the other hand, liquid precipitation may be supercooled (i.e., below freezing) as it falls through a sub-freezing layer, but if the ground temperature is greater than 0°C, it will not freeze upon contact with the surface, and thus be characterized as rain. The character of sleet is largely unchanged by ground temperature, except for the fact that ice pellets will melt if the ground temperature is greater than 0°C. Freezing rain and freezing drizzle will have little impact on ground temperatures when they reach the surface. In other words, ground temperatures may decrease slightly given the colder temperature of the droplets, but no melting, evaporation, or sublimation will occur to cool the temperature enough to change precipitation type. In fact, freezing rain and freezing drizzle often slightly raise the temperature of a sub-freezing ground surface by the phase-chang