Chapter 3: Convective Weather PDF

AFH15-101 5 NOVEMBER 2019 181 Chapter 3 CONVECTIVE WEATHER 3.1. Thunderstorms. Thunderstorm-produced severe weather consists of a combination of tornadoes, hail, strong winds, lightning, and heavy rainfall. There are three basic types of thunderstorms: single cell, multi-cell, and supercell. While u...

AFH15-101 5 NOVEMBER 2019 181 Chapter 3 CONVECTIVE WEATHER 3.1. Thunderstorms. Thunderstorm-produced severe weather consists of a combination of tornadoes, hail, strong winds, lightning, and heavy rainfall. There are three basic types of thunderstorms: single cell, multi-cell, and supercell. While upward vertical motions and instability of an air mass determine whether thunderstorms will occur, wind shear strongly influences the type of thunderstorms to expect. In general, the greater the shear, the more likely the convection will be sustained. Each type of storm can be identified by a distinctive hodograph pattern, which is a visual depiction of the wind shear. 3.1.1. Single cell thunderstorms. Single cell storms are short-lived (30 to 60 minutes) cells with one updraft that rises rapidly through the troposphere. Precipitation begins at the mature stage, in a single downdraft. When the downdraft reaches the surface, it cuts off the updraft and the storm dissipates. Figure 3.1 shows a typical hodograph for a single-cell storm. Single cell storm characteristics include weak vertical and horizontal wind shears, a random shear profile on the hodograph, and storm motion with the mean wind pattern in the lowest 5-7 km of the atmosphere. Severe weather does not normally occur with single cell storms, but may be possible in stronger and longer-duration cells. High winds and hail are possible, but shortlived, and tornadoes are rare. Figure 3.1. Single cell thunderstorm hodograph. 3.1.2. Multi-cell thunderstorms. Multicellular storms are clusters of short-lived single-cell storms; each cell generates a cold outflow that can form a gust front. Convergence along these boundaries causes new cells to develop every 5-15 minutes in the convergent zone. These storms are longer in duration than single cell storms, since they typically regenerate along the gust front. Figure 3.2 shows a typical hodograph for a multicellular storm. Multi-cell storms have a straight-line or unidirectional shear profile, and show strong directional shear in the lower levels with strong speed shear aloft. Individual cell motions coincide with the mean 182 AFH15-101 5 NOVEMBER 2019 wind, and storm clusters will move in the direction of the gust front and to the right of the mean wind. Severe weather is possible with multi-cell storms, including flash flooding from slowmoving cells, large hail near downdraft centers, and weak, short-duration tornadoes along gust fronts near updraft centers. Figure 3.2. Multi-cell thunderstorm hodograph. 3.1.3. Supercell thunderstorms. Supercell thunderstorms consist of a rotating updraft, a forward- flanking downdraft that forms the gust front, and a rear-flanking downdraft. Supercells may exist for several hours, and are a frequent producer of severe weather. They are characterized by wind speeds increasing with height, and a curved shear profile in lower levels, becoming straight-line above 3 km (Figure 3.3 shows a typical hodograph for a supercell.) They will have at least 70° of directional shear in the first 3 km of the atmosphere, and the shear vector veers with height in the low levels, which produces rotation of the storm updraft. A cyclonically-curved hodograph is associated with cyclonically rotating cells that will move to the right of the mean low-level wind; anticyclonically-curved hodographs indicate storms moving to the left of the mean wind – these types of storms are notorious hail producers. There are three types of supercells: classic, high precipitation (HP), and low precipitation (LP). Figure 3.3. Supercell thunderstorm hodograph. 3.1.3.1. Classic supercell (Figure 3.4). Classic supercells are usually isolated from the main thunderstorm outbreak, and are identified by the classic “hook echo” in the low-level reflectivity pattern and bounded weak-echo region (BWER) aloft. Supercells are capable of producing several types of severe weather: golf ball size hail, wind gusts in excess of 50 knots (along the gust front and from microbursts in the rear-flanking downdraft), and tornadoes. AFH15-101 5 NOVEMBER 2019 183 Figure 3.4. Classic supercell. 3.1.3.2. High-precipitation (HP) supercells (Figure 3.5). HP supercells develop in deep, moist layers with high moisture values; they produce heavier rain than classic supercells and are not as isolated. Radar patterns associated with HP cells are more varied than the classical “hook”, and have the potential to evolve into bow echo configurations. HP supercells are capable of producing extremely heavy rain, high winds, hail, and tornadoes. Figure 3.5. High-precipitation supercell. 184 AFH15-101 5 NOVEMBER 2019 3.1.3.3. Low-precipitation (LP) supercells (Figure 3.6). These types of supercells produce smaller amounts of precipitation than other supercell types, and have a rather benign appearance on radar. Although smaller in diameter than classic supercell storms, they are still capable of producing severe weather such as large hail and tornadoes. Figure 3.6. Low-precipitation supercell. 3.1.4. Dry, wet, and hybrid microbursts. Microbursts are dynamically enhanced, concentrated downdrafts from thunderstorms that result in damaging surface winds with gusts of 50 knots or greater at the surface. They usually occur in the rear-flanking downdraft region of supercell storms, and may also be found behind the gust front. These events are not restricted to large supercell storms, however; they can come from innocuous-looking, high-based rain clouds (dry microbursts), from single and multicellular pulse storms (wet microbursts), or from hybrid microbursts that combine dry and wet characteristics. The microburst type depends on the type of environment in which the storm forms; Figure 3.7 portrays typical atmospheric profiles for dry, wet, and hybrid microbursts. Figure 3.7. Microburst atmospheric profiles (Dry, Wet & Hybrid). AFH15-101 5 NOVEMBER 2019 185 3.1.5. Derechos. Derechos are straight-line wind events that originate from severe convective storms; there are two main variants of this event. The first type of derecho is a rapidlypropagating segment of an extensive squall line, usually associated with a strong, migratory low-pressure system occurring in the late winter or early spring. The second type develops in association with a relatively weak frontal system in a moisture-rich environment, showing characteristics of both squall lines and non-linear types of mesoscale convective systems (MCS) – these are usually late spring or summertime events. 3.1.6. Thunderstorm hazards – severe convective winds. The following are key components that must be examined when characterizing an environment conducive to convective winds: 3.1.6.1. Downdraft Convective Available Potential Energy (DCAPE). High downdraft instability is one of the key ingredients for severe convective winds. The amount of downdraft instability is diagnosed using a parameter called DCAPE; it’s a measure of the amount of energy available for air parcel descent through the atmosphere. As DCAPE increases past 800 J kg-1, the probability of strong downdrafts and severe convective winds increases. Values of DCAPE are increased by low-level warm air advection, low-level moisture advection, mid-level cold air advection, ans mid-level dry air advection (evaporative cooling increases negative buoyancy, increasing downdraft strength). Therefore, environments with steep low- and mid-level lapse rates (decreasing temperature) and dry mid-level air will have high values of DCAPE. If thunderstorms develop in such environments, strong downdrafts and severe convective winds are likely. The magnitude of DCAPE can also be diagnosed by examining the difference between the equivalent potential temperature (theta-e) in the low-levels and in the mid-levels. Theta-e is the temperature an air parcel would have if it was expanded (lifted) dry adiabatically to its LCL and then lifted moist adiabatically to the top of the atmosphere where all moisture condenses out of it. The parcel is then compressed (lowered) dry adiabatically to 1000 mb. High theta-e in the low-levels and low theta-e aloft (usually between 400 mb and 600 mb) is indicative of high downdraft instability. 3.1.6.2. Low and mid-level moisture profile. Convective winds may occur if the low-levels are moist and the mid-levels are dry, or if the low-levels are dry and the mid-levels are moist. They will not occur if both the low- and mid-levels are dry, because thunderstorms are unlikely to develop at all. They may occur if both the mid- and low-levels are moist due to precipitation loading, but are rare in this type of environment. Dry air aloft, characterized by relative humidity less than 50% allows much of the moisture that is lifted into the dry layer to evaporate. Evaporation is a thermodynamic process that cools the ambient air, increasing its density and decreasing the buoyancy. Dry air aloft is best characterized by examining values of equivalent or wet-bulb potential temperature between 600 mb and 400 mb. Both of these values account for the moisture and temperature characteristics of an air parcel, and the colder they are, the more negatively buoyant the air. In fact, downdrafts are believed to originate at the level of minimal wet-bulb or equivalent potential temperature aloft. If the mid-levels are moist (RH greater than 75%) and the lowlevels are dry (RH less than 75%), high-based thunderstorms may develop that also pose a convective wind threat. High-based precipitation may develop in the mid-levels. As it falls into the dry air below 700 mb, most of it will evaporate, resulting in additional cooling of the downdraft. In most cases like this, the downdraft is also sustained by steep low- and mid-level lapse rates during descent, which promote negative buoyancy. 186 AFH15-101 5 NOVEMBER 2019 3.1.6.3. Low-level vertical wind shear (VWS) profile. A strong VWS profile enhances the potential that strong winds aloft will be able to come to the surface, as long as the strongest wind in the surface to 6 km above ground level (AGL) layer is at least 20 knots. This additional mass increases negative buoyancy-driven downdraft magnitude. Steep low- and mid-level lapse rates allow winds from aloft to come to the surface, and this probability increases if the VWS profile (especially in the low-levels) is relative unidirectional, or inphase. The low-level VWS profile will also control the direction of thunderstorm propagation and influence the speed of the associated thunderstorm gust front. Stronger gusts are likely with fast-moving thunderstorms because of the propagation component and associated gust front speed. Most convective wind events are caused by organized thunderstorm structures in environments of sufficient VWS that also have favorable thermodynamic profiles. 3.1.6.4. Low-level lapse rates. Steep low-level lapse rates are not critical for convective winds, especially wet microbursts, but they do increase the negative buoyancy necessary for momentum transfer from aloft to the surface and for associated downbursts. Steep lapse rates are an indication of cold air aloft and warm air at the surface. Cold air on tap of warm air creates instability for both upward vertical motions and downward vertical motions. Lapse rates exceeding 8°C km-1, and especially approaching dry adiabatic, are optimal for convective wind development. Convective winds can also occur with lapse rates as shallow as 5°C km-1 if the low-levels are very moist and conducive to wet microbursts or generic downbursts due to precipitation loading, evaporation, and melting. Dry microbursts require very steep low-level lapse rates (usually dry adiabatic) for downdraft maintenance during descent, because the precipitation loading component is lacking throughout the downdraft’s evolution, except at the onset. In addition, downdrafts in a dry adiabatic environment tend to warm rapidly via compression. 3.1.6.5. Height of minimum wet bulb potential temperature aloft. Downdrafts are theorized to originate at the level of minimum wet bulb potential temperature (or equivalent potential temperature) aloft; the altitude of this level plays a role in the speed of descending air as it reaches the surface. In general, downdrafts that originate above 18,000 feet will have too much time to warm via compression during their descent, especially if the downdrafts are unsaturated, and resultant surface wind gusts will be weaker. Similarly, if downdrafts originate below 12,000 feet, the downdraft may not have enough time to accelerate before reaching the surface, rendering it weaker. The optimal layer for downdraft origination is between approximately 12,000 feet and 18,000 feet AGL (plus or minus 2,000 feet), or between 700 mb and 500 mb. Downdrafts originating in this layer experience an optimal balance of compressional warming vs. mass acceleration. 3.1.7. Thunderstorm hazards – hail. The following key parameters must be examined when characterizing an environment conducive to hail formation: 3.1.7.1. Mid-level lapse rates. For locations at mean sea level (MSL), lapse rates between 700 mb and 500 mb should be at least 6°C km-1, and optimally steeper than that. Locations at higher elevations should use the 600 mb and 400 mb lapse rate. Steep mid-level lapse rates lead to increased instability in the hail growth zone, which is usually between -10°C and -30°C. In this region, ice crystals and supercooled water droplets coexist in great supply. Instability in this layer promotes the vertical motions (via positive buoyancy) necessary for ice crystal formation and growth into hailstones. In addition, the steeper the AFH15-101 5 NOVEMBER 2019 187 mid-level lapse rates, the colder it is aloft and the easier it will be for hailstones to grow via riming of supercooled water droplets. Steep mid-level lapse rates also increase the depth of the hail growth zone, providing a greater area where hail growth is favored. 3.1.7.2. Wet bulb zero (WBZ) height. The WBZ zero height is the height at which the wet bulb temperature is 0°C. The WBZ height is a good proxy for the lowest level at which hail growth is possible; below the WBZ level, melting will inhibit hail growth. Most hail reaches the surface when the WBZ height is between 5000 and 11,000 feet AGL. The lower the WBZ level, the deeper the cloud depth where hail growth is possible. In addition, a low WBZ level means a shallower melting layer exists, so hailstones will have less time to melt before reaching the surface in such an environment. If the WBZ level is below 5000 feet AGL, the environment is likely too cold and stable for hail development. 3.1.7.3. Mid-level moisture profile. Dry air aloft, usually characterized by relative humidity less than 50% above 700 mb (or 600 mb for higher elevations), is favorable for severe hail. Dry air aloft aids in hail growth primarily by causing evaporative cooling that makes the upper-atmosphere more conducive to hail growth, and it also lowers the WBZ height. Dry air aloft is also an indicator of convective instability when the lower atmosphere is sufficiently moist. This will result in strong updrafts capable of supporting large hailstones. Convective instability is diagnosed by examining the profile of equivalent potential temperature, or theta-e. A layer where theta-e decreases with height is convectively unstable and will be more prone to hail production. Dry air aloft will usually be characterized by steep lapse rates as well, which further enhances instability and hail potential. 3.1.7.4. Convective instability. High instability, generally characterized by Convective Available Potential Energy (CAPE) greater than 500 J kg-1, favors the production of severe hail. Strong updrafts are needed for hailstone production and to allow hailstones to remain suspended so they can grow via riming, accretion, and aggregation. An environment of high instability (high CAPE) can support such strong updrafts. The magnitude of maximum upward vertical motion is directly proportional to the amount of CAPE; the larger the CAPE, the stronger the UVM and resultant thunderstorm updrafts. In general, for ½” hailstones to reach the surface, at least 500 J kg-1 of CAPE is necessary. At least 700 J kg1 is necessary for ¾” hailstones, and at least 2,000 J kg-1 is needed for 2” hailstones. 3.1.7.5. Surface to six km VWS profile. Strong VWS, generally characterized by surface to six km wind shear exceeding 15 knots, supports hail formation. Strong VWS enhances hail potential by increasing updraft strength via mesoscale pressure perturbations that result in a strong vertical pressure gradient force. It also ventilates thunderstorm updrafts by tilting storms downshear, so the downdrafts that develop don’t fall into the updrafts. In other words, strong VWS keeps downdrafts away from updrafts, which allows hailstones to be suspended for a longer period and allows updrafts to maintain their strength. Large hail is possible in environments with CAPE as low as 500 J kg-1 that have strong VWS. In fact, hail is more likely in low CAPE/high VWS environments than in those characterized by high CAPE and low VWS. For maximum severe hail potential, both high CAPE and strong VWS are necessary. This is most commonly the case during the spring months in the mid-latitudes, especially late April and May. 188 AFH15-101 5 NOVEMBER 2019 3.1.7.6. Low-level moisture. Low-level moisture is critical for hail production; moisture must be present to be transported to the LCL, where hailstones originate as water droplets, and then into the -10°C to -30°C hail growth zone, where graupel and eventually hailstones form. The best variable for identification of deep low-level moisture is Precipitable Water (PW); this is the depth of water that would accumulate if all the water in the atmosphere fell out as precipitation. For severe hail formation, PW should be less than 1.5” but greater than 0.5”. While not a critical problem for severe hail development, higher amounts of PW (greater than 1.5”) increase the likelihood that any updrafts that are able to develop will become loaded with liquid water, therefore decreasing their strength and also the likelihood that they can support severe hail. The largest hailstones are observed in environments that have high CAPE, strong VWS, and low PW. However, if the environment is too dry as indicated by PW values less than 0.5”, the positive buoyancy needed for updraft development and maintenance will not be achieved, reducing hail potential. 3.1.7.7. Surface elevation. High elevation locations, especially those where the surface is above 900 mb (approximately 4000 feet), are most favored for large hail. Because temperature generally decreases with increasing height in the troposphere, locations with higher elevations will have easier access to cold air aloft. This means that, on average, the WBZ level is closer to the surface. Consequently, the melting layer below the WBZ level is shallower, giving hailstones that develop above the WBZ level a shorter depth in which to melt. In addition, because wind speeds generally increase with height in the troposphere, higher elevations are able to tap into stronger winds aloft, which enhance the surface to six km VWS profile. 3.1.8. Thunderstorm hazards – heavy rainfall. The following key parameters must be examined when determining if an environment is favorable for heavy rain: 3.1.8.1. Precipitable Water (PW). The higher the PW, the more water vapor that is available to condense and eventually turn into precipitation. Generally, precipitable water of at least 1 inch (or greater than 100% of normal) is necessary for heavy rain. Values of 1.5 inches (or greater than 150% of normal) characterize a very favorable heavy rain environment. 3.1.8.2. Relative humidity in the lowest 200 mb of the atmosphere. The lifting of a deep layer of moist air results in more efficient precipitation production than the lifting of a shallow layer of moist air. Therefore, RH of at least 70% in the lowest 200 mb of the atmosphere is generally necessary for heavy rain. An even deeper layer of high RH (up to 500 mb) further increases precipitation efficiency by reducing entrainment of dry midlevel air. 3.1.8.3. Surface dew points. The dew point is the temperature to which air must be cooled at constant pressure for water vapor to condense into liquid. Because most water vapor is confined to the lowest 200 mb of the atmosphere, near moisture sources (oceans, lakes, rivers), the surface dew point is an important proxy for the amount of water vapor available for precipitation development. As the surface dew point exceeds 55°F and approaches 60°F and higher, the amount of low-level water vapor available for precipitation production increases substantially, and so does the chance of heavy rainfall. Surface dew points are regularly in excess of 65°F during the late spring and summer months in the lower and mid-latitudes, which is when many heavy rain events occur. AFH15-101 5 NOVEMBER 2019 189 3.1.8.4. Moisture convergence. Moisture convergence is the advection of water vapor (moisture) against a convergent boundary such as a front, coastline, or complex terrain. Moisture speed convergence can also occur as stronger winds meet up with slower winds, causing a local increase in dew points and relative humidity. Moisture convergence acts as a source of low-level lift for precipitation production, and also increases dew points and relative humidity as more water vapor is brought into an area and “pools” along a boundary. Greater horizontal and vertical depth of moisture inflow and convergence increases heavy rainfall potential. Of particular importance is persistent southerly inflow and convergence within the left exit region of a low-level jet along a quasi-stationary low-level frontal boundary. This can signify the potential for several inches of rainfall. Once convection has developed, continued moisture inflow and convergence can maintain the convection, even in the absence of other forcing mechanisms. 3.1.8.5. K index. The K index accounts for both stability and the presence or lack of deep moisture available for precipitation and thunderstorm development; it’s given by: KI = (T850 – T500) + (Td850 – DD700). Where T850 is the temperature at 850 mb (°C), T500 is the 500 mb temperature (°C), Td850 is the 850 mb dew point (°C), and DD700 is the dew point depression at 700 mb in °C. High instability and deep moisture result in large K index numbers; values of at least 25, and especially over 30, are indicative of heavy rainfall potential. 3.1.8.6. Surface to six km wind shear. Thunderstorms that develop in a low vertical wind shear environment move slowly and may dump heavy rain over the same locations for several hours. Heavy rainfall is most favored when surface to 6 km wind shear is less than 15 knots. 3.1.8.7. Winds directed from a moisture source in the lowest 200 mb above the surface. Low-level winds (surface to 850 mb) directed from a moisture source are another key ingredient for heavy rainfall. Winds directed from a moisture source continuously advect moisture into a region, and if thunderstorms are ongoing, they serve as moist inflow for thunderstorm maintenance as other moisture precipitates out. This is why much more precipitation can fall during a strong thunderstorm than indicated by the PW value. Moist inflow continuously replenishes the lower atmosphere and maintains or increases the current value of PW. 3.1.8.8. Equivalent potential temperature (theta-e) advection. Theta-e is the temperature air would have if it was lifted dry adiabatically to its LCL, and then lifted moist adiabatically from the LCL to the top of the atmosphere, allowing all water vapor to condense out, then lowered dry adiabatically to 1000 mb. Theta-e is an excellent diagnostic of both the temperature and moisture content of air. Warm, moist air will have higher thetae values than cold, dry air. Therefore, strong positive theta-e advection into an area of responsibility signifies an increased potential for heavy rain. Heavy rain potential is further enhanced if the low-level, high theta-e air impinges on a boundary such as a front or complex terrain, leading to moisture convergence. 190 AFH15-101 5 NOVEMBER 2019 3.1.8.9. Thickness diffluence. Organized convection, accompanied by heavy rainfall, can occur in or near a region where 1000-500 mb thickness isopleths are diffluent. Thickness diffluence implies low-level convergence, upper-level divergence, or both, and is therefore favorable for convective development. Thickness diffluence is usually located along or near the southern edge of the mid-tropospheric westerlies. 3.1.8.10. Jet streams. There are four regions of jet-enhanced divergence that are favorable for convective development and associated heavy rainfall: the right entrance region of a jet streak, the left exit region of a jet streak, the exit region of a jet streak approaching the top of a ridge axis, and the anticyclonic shear axis to the right of a jet core. Of particular importance is a coupled jet streak, which causes enhanced lift and thunderstorm potential. Additionally, coupling between an upper-level jet streak and the low-level jet can enhance thunderstorm potential by increasing low-level convergence and lift toward an area of upper-level divergence. Heavy rainfall is also likely near the nose and/or the left side of the low-level jet axis, where speed convergence, confluent flow, frontogenesis, and lift are maximized. 3.1.9. Thunderstorm hazards – lightning. The following key parameters must be analyzed when determining if the environment is favorable for lightning production: 3.1.9.1. K index. As discussed in the previous section, the K index accounts for both stability and the presence or lack of deep moisture available for precipitation and thunderstorm development; high instability and deep moisture result in large values, indicative of lightning potential. Deep low-level moisture is required for the eventual creation of ice crystals, graupel, and supercooled water droplets aloft. Supercooled water droplets collide with ice crystals to form graupel. When graupel and ice crystals collide, a charge transfer occurs. Strong vertical motions, potentially indicated by a high K index value, drive the lighter ice crystals toward the top of the cumulonimbus cloud, while the heavier graupel particles fall to the middle or lower portions, and a charge separation is created. Instability (and a source of lift) is needed for the aforementioned moisture to be transported to higher heights, condense/coalesce, and turn into supercooled water droplets, ice crystals, and graupel. K index values greater than 30 indicate a strong likelihood of efficient lightning production. 3.1.9.2. Lifted index (LI). The LI is an indicator of stability, particularly at mid-levels (i.e., the 700-500 mb layer). The equation for the LI is: LI = Te500 – Tp500. Where Te500 is the temperature of the environment at 500 mb and Tp500 is the temperature of an air parcel lifted to 500 mb from near the surface. LI values less than or equal to -1°C (less than or equal to +1°C for locations above 1,000 meters) indicate the presence of mid-level instability, which supports upward vertical motion within the layer where lightning production is favored, generally between -10°C (lower to middle portion of a cumulonimbus cloud) and -25°C (upper portion of a cumulonimbus cloud). Further, lift in the 0°C to -20°C layer favors the creation of supercooled water droplets, ice crystals, and graupel. AFH15-101 5 NOVEMBER 2019 191 3.1.9.3. Showalter Stability Index (SSI). The SSI is calculated by lifting a parcel of air dry adiabatically from 850 mb to its lifting condensation level, then moist adiabatically to 500 mb. The 500 mb parcel temperature is then subtracted from the 500 mb environmental temperature; if the SSI is greater than zero, the environment is stable. Increasingly negative values indicate greater instability and thunderstorm potential. Note, however, that lowlevel moisture must extend to at least 850 mb for the SSI to be representative. Several thunderstorm studies showed that 81% of days without lightning had SSI values greater than 0, while 90% of lightning days had SSI values less than 0. Using the predictor of greater than/less than 0, the SSI accurately predicted whether lightning would occur 96% of the time. 3.1.9.4. Convective Available Potential Energy (CAPE) between 0° and -20°C. Another key parameter for lightning production is at least 200 J kg-1 of CAPE in the 0°C to -20°C layer. Lightning production becomes especially likely when CAPE exceeds 600 J kg-1. CAPE is a measure of the energy available for air parcel ascent from the level of free convection to the equilibrium level. However, large CAPE in the 0°C to -20°C layer is especially important for lightning formation because this is the layer where ample upward vertical motion is needed to allow the microphysical processes to take place which create supercooled water droplets, ice crystals, and graupel. The larger the CAPE in this layer, the stronger the ascent, and the more efficient the microphysical processes become. 3.1.9.5. Convergence in the planetary boundary layer (PBL). Mechanical or moisture convergence in the PBL (the lowest 2 km of the troposphere, on average) favors lightning production. Convergence in this layer lifts water vapor upward where it can eventually form into supercooled water droplets, ice crystals, and graupel. 3.1.9.6. Capping in the PBL. A weak cap to upward vertical motion, represented by a Lid Strength Index (LSI) less than 6°C and/or Convective Inhibition (CIN) greater than -100 J kg-1, is favorable for lightning. Weak capping does not overly suppress upward vertical motion, which allows moisture to be carried aloft for the development of supercooled water droplets, ice crystals, and graupel. 3.1.9.7. Relative humidity about the lifting condensation level (LCL). RH should generally be greater than or equal to 90% in the 75 mb layer above the LCL if thunderstorm development is to occur. (Note: For areas above 1000 meters, RH should be greater than or equal to 70% in the 50 mb layer above the LCL.) Air parcels that are saturated at the LCL become diluted by dry air as they ascend above the LCL when the relative humidity is too low; dry air decreases the buoyancy of the ascending parcels, and often causes their ascent to cease. 3.1.9.8. Precipitable Water. When precipitable water steadily increases over time (at least 0.1” in 6 hours) and/or exceeds 150% of normal, lightning becomes increasingly likely. Steadily increasing and/or high PW is indicative of increasing/substantial water vapor available for the eventual creation of supercooled water droplets, ice crystals, and graupel, all of which contribute to the thunderstorm electrification process. 3.1.10. Thunderstorm hazards – tornadoes. The following key parameters must be analyzed when determining if the environment is favorable for tornadoes: 192 AFH15-101 5 NOVEMBER 2019 3.1.10.1. Zero to three km wind profile. An environment conducive to tornadogenesis will be characterized by a veering (clockwise turning) wind profile in the lowest 3 km of the troposphere. Tornadoes can develop with veering or backing winds in the lower troposphere, and even with straight winds that have strong speed shear. However, in the Northern Hemisphere, the majority of tornadoes occur in environments where the lowlevel winds veer with height. A particularly favorable environment will feature surface winds from the southeast at greater than 15 knots. The veering winds create positive horizontal vorticity (spin) that can be tilted into the vertical by an updraft. 3.1.10.2. Zero to six km wind shear. A strong, deep layer of vertical wind shear is favorable for the development of tornadoes; a favorable shear profile is generally characterized by wind shear from the surface to six km of at least 30 knots, with values greater than 40 knots being optimal. This magnitude of deep layer vertical wind shear ventilates thunderstorm updrafts, separating them from negatively buoyant downdrafts. In this manner, the updraft of a supercell can be maintained for several hours. Additionally, substantial 0-6 km wind shear creates a strong vertical pressure gradient, which enhances updrafts already produced by mechanical lift (fronts, low-convergence, etc.) and positive buoyancy (sufficiently large CAPE). 3.1.10.3. Inflow layer wind speeds. If warm, moist winds exceeding 15 knots are ingested into a thunderstorm’s inflow layer, updraft strength will be enhanced and updrafts will be maintained. Further, strong inflow layer winds can enhance the low-level vertical wind shear profile, increasing tornado potential. Once inflow is cut off or becomes too weak, thunderstorm outflow from negatively buoyant and precipitation-loaded downdrafts tends to overwhelm ongoing thunderstorms, thereby eliminating the potential that they will spawn tornadoes. 3.1.10.4. Storm Relative Helicity (SRH). SRH estimates a thunderstorm’s potential to acquire a rotating updraft given an environmental vertical wind shear profile. Larger SRH values are indicative of a higher probability that rotating updrafts will develop in ongoing thunderstorms. SRH is not a “tornado predictor” by itself; it’s conditional upon thunderstorm development, and not useful unless thunderstorms are ongoing, since it’s calculated in a storm relative framework. Traditionally, 0-3 km SRH is used to evaluate storm type and rotation potential; values greater than 250 m2 s-2 suggest an increased threat of tornadoes with supercells, and 0-3 km SRH values exceeding 300 m2 s-2 are usually found in environments that produce violent tornadoes. 0-1 km SRH is an indicator of the potential for low-level rotation, and has been found to be useful for distinguishing between non-tornadic and tornadic supercells. At least 100 m2 s-2 of 0-1 km SRH is needed for tornadogenesis on most occasions. Frontal and outflow boundaries often focus and enhance SRH, leading to an increased probability of tornadogenesis. 3.1.10.5. CAPE. Tornadoes can occur within a very wide range of CAPE values, but most tornadoes occur when CAPE is between 1,500 J kg-1 and 4,000 J kg-1. Maximum updraft strength is directly proportional to the amount of CAPE, so the larger the CAPE, the stronger the resultant updraft that can develop to tilt horizontal vorticity into the vertical. AFH15-101 5 NOVEMBER 2019 193 3.1.10.6. Convective Inhibition (CIN). A weak cap to upward vertical motion, indicated by CIN greater than -100 J kg-1, is favorable for severe thunderstorm development because excess energy builds up below the cap. This energy is released if the cap breaks, leading to explosive thunderstorm development. However, if the cap is too strong, thunderstorms will have a difficult time developing in the first place and tornado potential will be greatly reduced. 3.1.10.7. Bulk Richardson Number (BRN). The BRN is a non-dimensional ratio of CAPE to vertical wind shear, and is used to characterize connective-storm types for various environments. Small CAPE and large shear values result in small BRN values (commonly found during the late autumn and winter months); thunderstorms may develop in this environment, but they are likely to be sheared apart. Conversely, large CAPE and small shear values result in large values of the BRN (commonly found during the summer months); thunderstorms developing in this environment are likely to be pulse storms that do not develop into supercells. Values of BRN between 10 and 45 are indicative of an optimal balance of CAPE and vertical wind shear for supercell and possible tornado development. 3.1.10.8. The lifting condensation level (LCL). The LCL is the level at which an unsaturated ascending air parcel becomes saturated; it’s an excellent proxy of cumuliform cloud bases. An LCL near the surface indicates the presence of substantial boundary layer moisture, and thunderstorm updrafts are therefore less likely to entrain dry air. This increases positive buoyancy and the potential for tornadogenesis. Additionally, when the LCL is low, cold outflow is less likely to undercut the updraft and interfere with developing tornadic circulations. Most tornadoes occur when the LCL is below 1,500 meters, and significant (EF2+) tornadoes almost always occur when the LCL is 1,000 meters or lower. 3.2. Synoptic Patterns. There are several basic synoptic weather patterns that can produce severe weather in the mid-latitudes; areas for likely thunderstorm development often experience some combination of mid-level jets or wind shear, dry-air intrusions between 850 mb and 700 mb, and low-level moisture gradients. These parameters are proven severe thunderstorm triggering mechanisms, and can help identify where severe thunderstorm outbreaks will occur in each of the synoptic patterns. Mid-level jets (wind speed and shear maxima between 700-500 mb) can indicate areas of thunderstorm and tornado development. Dry-air intrusions at 700 mb are a major triggering mechanism for tornadoes, and can be used to pinpoint areas of potential severe thunderstorm development. Dry-air intrusions are difficult to identify by a particular temperature/dew-point spread or relative humidity, since the values vary widely from case to case. They can often be identified by looking at the intensity with which drier air is being forced into the moist air. Most severe thunderstorm outbreaks are associated with strong low-level (below 700 mb) moisture gradients; the gradient axes are generally located on the windward side of the outbreak area. The intensity of the storm is proportional to the tightness of the moisture gradient along the wind component from dry to moist air. When the 850 mb or 925 mb product is not representative of moisture below 700 mb, the moisture gradient can often be determined from satellite imagery and model-generated vertical cross sections. Table 3.1 shows an empirical relationship between various values of these parameters and the potential for severe thunderstorm development. 194 AFH15-101 5 NOVEMBER 2019 Table 3.1. Severe thunderstorm development potential. Parameters Mid-level jet speed Weak 35 knots Moderate 35-50 knots Horizontal shear 15 knots/90 NM 15-30 knots/90 NM Less than 20° 20°-40° Greater than 40° Less than 13°C Less than 8°C 13°C-18°C 8°-12°C Greater than 18°C Greater than 12°C Wind crossing the axis of 700 mb dry intrusions and moisture boundaries Surface dew point 850 mb dew point Strong Greater than 50 knots Greater than 30 knots/90 NM 3.2.1. Classic Synoptic Convective Weather Patterns. Identifying severe synoptic patterns is essential to identifying areas of potentially severe thunderstorms; successful severe weather forecasting is dependent on the ability to analyze, coordinate, and assess the relative values of a multitude of meteorological variables and mentally integrate and project these variables three-dimensionally in space and time. Identification of severe synoptic patterns saves time and allows a focused effort on the threat area. In the severe weather patterns outlined below, the parameters are morning depictions (12Z for CONUS), while the outbreak areas are depicted at the time of occurrence, which may later in the day; advection of severe weather parameters must be taken into account. 3.2.1.1. Type A synoptic pattern (dryline). With this type of synoptic pattern (Figure 3.8), thunderstorms initially form on the edge of a sharp moisture gradient. Storms tend to form rapidly in widespread, isolated clusters. 3.2.1.1.1. Characteristics. The type A pattern is defined by a well-established southwesterly 500 mb jet, a distinct surface-to-700 mb warm dry-air intrusion from the southwest, low-level confluence along the dry line, and low-level moisture advection from the south, ahead of the dry air. Convective development with this pattern is characterized by extremely rapid growth (15-30 minutes) from inception to maturity, with almost immediate production of large hail, damaging winds, and tornadoes. 3.2.1.1.2. Initial outbreak area. Severe storm formation is usually confined to the edge of the dry air at 850 mb and 700 mb, and the convergence area between the moist and dry air (the area of maximum moisture gradient). These storms will form rapidly, in isolated clusters, along the leading edge of the dry intrusion. Sharp, well-defined squall lines aren’t common with this pattern. AFH15-101 5 NOVEMBER 2019 195 Figure 3.8. Type A severe weather synoptic pattern - dryline. 3.2.1.1.3. Severe weather area. In a type A pattern, severe weather typically extends up to 200 miles to the right of the 500 mb jet, and from the area of maximum low-level convergence to the area of rapidly decreasing moisture along the moisture gradient. The most violent storms usually form where the jet meets the moist/dry air convergence area. A secondary outbreak area can occur along and 150 miles to the right of the 500 mb horizontal speed shear zone; it extends from the maximum low-level convergence area to the point where low level moisture decreases on the dry side of the moisture gradient. 3.2.1.1.4. Trigger mechanisms. Type A storm development can be triggered by several factors, such as diurnal heating, passage of an upper-level jet max, low level intrusions of warm, moist air east of the dryline, or mid-level dry air moving into a moist region. 3.2.1.1.5. Timing. In type A situations, look for thunderstorms to develop at the time of maximum heating, or up to 6 hours afterwards. Under normal circumstances, convection is usually capped by an inversion until the convective temperature is reached. Once convective activity has started, expect it continue for at least 6-8 hours, and possibly longer. The convective activity may last until the moist and dry air are completely mixed, changing the airmass’s stability. 3.2.1.2. Type B synoptic pattern (frontal). The Type B pattern (Figure 3.9) is defined by prefrontal squall lines with one or more mesoscale lows. These squall lines form at the intersection of the low-level jet and the upper-level jet. The lows often form in the area of the intersection of the low-level jet and the warm front and are frequently accompanied by tornadic outbreaks. 196 AFH15-101 5 NOVEMBER 2019 Figure 3.9. Type B severe weather synoptic pattern - frontal. 3.2.1.2.1. Characteristics. Type B systems are characterized by a well-defined 500 mb jet stream, a well-defined dry air intrusion between the surface and 700 mb, a strong unstable wave with associated warm and cold fronts, and a low-level jet, which is instrumental in transporting warm, moist air from the south. Most type B systems have frontal and/or pre-frontal squall lines, with strong cold-air advection behind the front. There will also be cool, moist air present along the 500 mb and 700 mb trough axes – the axes will lie to the immediate west of the threat area. Low and mid-level confluence between low-level warm air and mid-level cooler air will also be present. 3.2.1.2.2. Severe weather. The severe weather occurring with the type B pattern is associated with strong cold air advection and strong cold fronts. This type of system can occur at any time of the year, but it’s most frequent in the spring. As cold air moves into the threat area, it collides with warm, moist air advecting from the south. This collision of contrasting air masses leads to strong thunderstorms. 3.2.1.2.3. Severe weather areas. Several events occur in the initial severe weather area, beginning with the development of mesoscale lows at the intersection of the low-level jet and the warm front; as the meso-low develops, upward vertical motion in the area is increased. The location of severe weather depends on the speed of the cold front, coupled with the speed of the dry intrusion area. The best potential for severe weather is along and 150 miles to the right of the horizontal speed zone of the upper-level jet, but the area of concern can extend down to the leading edge of the dry air intrusion. The threat area does not extend into the dry air, as the absence of moisture decreases the chance of thunderstorm development. 3.2.1.2.4. Trigger mechanisms. The main trigger for type B systems is an approaching cold front, coupled with the dry air intrusion; the cold front provides lift, and the dry air decreases stability. Intersecting lines of discontinuity can also spur initial development; watch for intersecting squall lines or upper and lower-level jet streams, and the intersection of a low-level jet with a warm front. AFH15-101 5 NOVEMBER 2019 197 3.2.1.2.5. Timing. In this pattern, thunderstorms can occur anytime, and may last all day and night. Type B thunderstorms do not require diurnal heating, and as long as the airmass stays unstable, thunderstorms in a squall line can persist. 3.2.1.3. Type C synoptic pattern (overrunning). Type C patterns (Figure 3.10) generate severe weather conditions by overrunning – warm, moist air overrunning cold, dense air. If the warm air is sufficiently unstable, the lift over the cold air enhances upward motion and enhances the development of thunderstorms. 3.2.1.3.1. Characteristics. Type C patterns have an east-to-west oriented stationary front, with warm, moist overrunning tropical air. In addition, a west-southwest to westnorthwest upper-level jet or a strong 500 mb westerly horizontal wind speed shear zone will be present. There will also be a 700 mb dry intrusion advecting in from the southwest. 3.2.1.3.2. Severe weather. Damaging winds and large hail are possible in a type C pattern; tornadoes may occur when surface dew points are 50°F or higher; latent heat release at dew points that high may provide sufficient energy for tornado formation. 3.2.1.3.3. Severe weather areas. Scattered thunderstorms will develop on and north of the stationary front due to the overrunning; in the overrunning region, a squall line may form along the leading edge of the dry air intrusion and thunderstorms may reach severe levels. The severe threat area extends from approximately 50 miles west of the axis of maximum overrunning to the eastern edge of the overrunning. Figure 3.10. Type C severe weather synoptic pattern – overrunning. 3.2.1.3.4. Trigger mechanisms. The overrunning is the chief trigger for type C scenarios; maximum diurnal heating and a dry-air intrusion where thunderstorms are already occurring are several other triggers under this pattern. 198 AFH15-101 5 NOVEMBER 2019 3.2.1.3.5. Timing. Severe thunderstorm occurrence and duration depends on the onset time of dry air intrusion and maximum heating; severe weather continues until the dry air intrusion decreases or moves out. Severe activity can last up to 6 hours after maximum heating. 3.2.1.4. Type D synoptic pattern (cold core). Type D patterns (Figure 3.11) are noted for funnel clouds and large hail. Tornadoes are rare, but can occur with strong systems. The defining feature of a type D system is the cold core. 3.2.1.4.1. Characteristics. Type D systems have a deep, southerly upper-level jet, along with a deepening surface low, a 500 mb cold-core low, and cool dry air advection at all levels. There will also be a low-level jet advecting warm, moist air from the southsoutheast, under the cold air aloft. 3.2.1.4.2. Severe weather. Hail and funnel clouds may occur with these types of scenarios. Funnel clouds in the Type D Pattern are often referred to as “cold air” funnels, as their formation is caused by warm air moving under cold air aloft, which is associated with a cold core low at 500 mb. Figure 3.11. Type D severe weather synoptic pattern – cold core. 3.2.1.4.3. Severe weather areas. Thunderstorms will form in the area between the upper-level jet and the closed isotherm center at 500 mb. Severe weather may occur in the region bounded by the cold core low, the front edge of the dry air intrusion, the east-northeast limit of the underrunning warm air, and the area approximately 150 miles right of the upper-level jet. 3.2.1.4.4. Trigger mechanisms. There are two main triggers for type D systems; intense low-level confluence and decreasing stability due to the upper-level cold air moving over warm, moist air. 3.2.1.4.5. Timing. Severe weather typically occurs during or shortly after max heating, with a rapid decrease in intensity after sunset. AFH15-101 5 NOVEMBER 2019 199 3.2.1.5. Type E synoptic pattern (squall line). With type E systems (Figure 3.12), frontal or prefrontal squall lines are usually well defined; the squall lines may be fast or slow moving. In either case, severe storms develop rapidly. 3.2.1.5.1. Characteristics. Type E systems have a well-defined upper-level westerly jet, along with a well-defined dry air region bounded by a 700 mb warm sector. Low level convergence will be present, along with moderate-to-strong southerly low-level flow advecting warm, moist air over cooler, drier air. In addition to developing ahead of cold fronts, squall lines associated with the type E pattern also develop ahead of warm and occluded fronts. Squall lines may form as a line, or may organize into a line from a cluster of cells; the component of low-level winds shear perpendicular to the line orientation is the most critical factor for squall line structure and evolution. A typical squall line life cycle is to evolve from a narrow band of intense convective cells to a broader, weaker system. The timing of this evolution is strongly dependent on the magnitude of the low-level vertical wind shear; stronger shear leads to longer-lived squall line systems. Figure 3.12. Type E severe weather synoptic pattern – squall line. 3.2.1.5.2. Severe weather areas. Severe weather may develop along and south of the upper-level jet but north of the 850 mb warm front. The west-east boundary is from the 700 mb cold front to the area of increasing stability. Thunderstorms form in the overrunning warm air between the 850 mb warm front and the upper-level jet axis, where the 700 mb dry air intrusion meets the frontal lifting of the warm, moist air in the low-levels, and the strong 500 mb cold air advection. A secondary threat area exists where the 700 mb dry air intrusion extends south of the 850 mb warm front. Thunderstorms can develop along the 500 mb horizontal speed shear zone and along transitory, active squall lines. 3.2.1.5.3. Trigger mechanisms. The main drivers for severe development in type E systems are frontal lifting of warm, moist and unstable air, a 700 mb dry air intrusion, diurnal heating, and cold air advection at 500 mb. 200 AFH15-101 5 NOVEMBER 2019 3.2.1.5.4. Timing. Thunderstorms develop with the onset of 500 mb cold air advection into the severe outbreak area; maximum severe activity occurs from the time of maximum heating to a few hours after sunset. Severe storms may continue until midnight, or until the airmass becomes more stable. 3.3. Convective Weather Tools. The thermal instability of an air parcel can be expressed as a numerical value using a wide variety of stability indices; these tools are aids for determining severe weather potential, and should not be used as the sole basis for making a thunderstorm forecast. 3.3.1. Stability Indices. Table 3.2 lists various general thunderstorm indices and threshold values, and Table 3.3 shows indices and threshold values for severe weather potential. Table 3.4 and Table 3.5 show various tornado indicators. Thresholds vary from location to location, so closely monitor these indices to discover the best value for local use and adjust accordingly; the best way to evaluate a threshold is to keep a continuous record of their effectiveness. Regional values are provided where available. Each index is described in further detail below. 3.3.1.1. Convective Available Potential Energy (CAPE). CAPE is a measure of the convective instability of the atmosphere and thus, the potential for thunderstorms. CAPE values are not a direct indicator of severe weather; they should be used in conjunction with helicity (a measure of the rotation potential of a column of air). Depending on helicity, severe thunderstorms and tornadoes can occur under a wide range of CAPE values. AFH15-101 5 NOVEMBER 2019 201 Table 3.2. General thunderstorm (instability) indicators. 3.3.1.2. Bulk Richardson Number (BRN). Discussed in detail earlier in the chapter, the BRN is a better indicator of storm type than of storm severity or storm rotation; it’s useful in differentiating between weak, multi-cellular storms (non-severe) and supercell-storm (severe) types. This index is a measure of turbulent energy (a ratio of buoyancy to vertical wind shear) in a column of air to enhance or hinder convective activity. The BRN is most accurate when the CAPE index is between 1500 and 3500 J/kg; when CAPE is less than 1000 J/kg and accompanied by moderate wind shear, the BRN may indicate supercells, but the lack of buoyancy is likely to inhibit severe weather occurrence. If CAPE is greater than 3500 J/kg with a moderate wind shear environment, BRN values may suggest multi-cell storms (non-severe storms), but the buoyant energy will be sufficient to produce tornadoes and large hail. BRN values between 10 and 45 are indicative of supercell and possible tornado development. 202 Table 3.3. Severe thunderstorm indicators. Table 3.4. Tornado indicators. AFH15-101 5 NOVEMBER 2019 AFH15-101 5 NOVEMBER 2019 203 Table 3.5. Tornado forecasting tools. 3.3.1.3. Cross Totals (CT). The CT index compares low-level moisture and upper-level temperature; it’s optimized for thunderstorm coverage and severity east of the Rockies and along the Gulf Coast. The CT value is contingent on the low-level moisture band being at 850 mb and the cold air pocket at 500 mb; if the moisture and cold air are centered slightly above or below these levels, CT values will not be a reliable indicator of thunderstorm coverage or severity. 3.3.1.4. Dynamic Index. This index is designed for airmass thunderstorms; positive values indicate stability, and negative numbers indicate a conditionally unstable air mass. A triggering mechanism is needed for thunderstorms to occur when conditionally unstable; diurnal heating is usually enough to trigger the convection. 3.3.1.5. Energy/Helicity Index (EHI). The EHI should only be used if strong thunderstorms are forecast. EHI is a combination of CAPE and Storm Relative Helicity (S-RH), which measures the contribution of convective instability of the atmosphere and shear vorticity to the potential for tornado formation. Strong-to-violent tornadoes are associated with a wide range of CAPE values: large CAPE values combined with low wind shear and low CAPE values combined with high wind shear are both capable of producing conditions favorable for the development of tornadoes. 3.3.1.6. Fawbush-Miller Stability Index (FMI). This index is similar to the Showalter Stability Index, except it emphasizes the low-level (surface) moisture rather than the 850 mb moisture. Only use the FMI when the Showalter appears to be misrepresenting the lowlevel moisture. 3.3.1.7. GSI Index. This index was developed for use in the central Mediterranean, using the following procedure: 3.3.1.7.1. Obtain the minimum temperature/dew point spread (°C) between 650 mb and 750 mb. 3.3.1.7.2. Obtain the average wet-bulb temperature in the lowest 100 mb by the equal area method. From this point, follow the saturation adiabat to the 500 mb level. Subtract the temperature where the saturation adiabat crosses the 500 mb level from the observed 500 mb temperature (°C). 3.3.1.7.3. Add the values from Step 1 and Step 2 above to calculate GSI. 204 AFH15-101 5 NOVEMBER 2019 3.3.1.8. K index. As illustrated earlier in this chapter, the K index is primarily used for forecasting heavy rain and lightning potential; it is not an indicator of severe weather. It works best in the summer east of the Rockies in maritime-tropical (mT) air masses and in any tropical region. It has limited use in overrunning situations and in mountainous regions. 3.3.1.9. KO Index. The KO index, created by the German Weather Bureau, is sensitive to moisture and works best for cool moist climates (i.e., Europe, Pacific Northwest). The KO Index‘s drawback is its complexity. The KO equation is: KO = (Qe500 + Qe700)/2 – (Qe850 + Qe1000)/2. Where Qe is the equivalent potential temperature at a given level. To find Qe, first find the lifting condensation level (LCL) for the given pressure level. Continue up the moist adiabat until all moisture is removed from the parcel. This occurs at the level where the moist and dry adiabats become parallel. From there, continue up the dry adiabat to the top edge of the chart. There, read Qe directly. Do this for each of the four pressure levels in the equation and plug into the equation. The result is the KO index. 3.3.1.10. Lifted Index (LI). As discussed earlier in this chapter, the LI is useful for lightning potential determination; it can be used successfully at most locations, since it contains a good representation of the low-level moisture. This index counters deficiencies in the Showalter Index when low-level moisture and/or inversions are present. However, it fails to consider cold air above 500 mb. Threshold values are generally lower than the Showalter Index. 3.3.1.11. Modified Lifted Index (MLI). The MLI considers the destabilizing effects of cold air aloft, which the LI fails to take into account. It works well as a severe thunderstorm indicator in Europe, and has also been used with success in the CONUS. It gives poor results when the -20°C level is above 500 mb (too warm) or below the LCL (too cold). 3.3.1.12. S Index. The German Military Geophysical Office developed the S index as a variation of the Total Totals (TT) index. The S Index adds moisture available at 700 mb to a variable parameter based on the Vertical Totals Index (VT). The addition of 700 mb moisture tailors this index for sections of Europe since low-level heating is usually less intense in parts of Europe than it is in the States, and 700-mb moisture is a good predictor of thunderstorm development there. The S Index is useful from April to September. It can be computed from the equation shown in Table 3.6. Table 3.6. S-Index calculation. S = TT – (700T – 700Td) – A where A is defined as follows: VT Greater than 25 Greater than 22 but less than 25 Less than 22 A 0 2 6 3.3.1.13. Severe Weather Threat Index (SWEAT). The SWEAT index is designed to predict severe storms and tornadoes, rather than ordinary thunderstorms. High SWEAT values do not necessarily mean that severe weather will occur, since it doesn’t consider triggering mechanisms. High SWEAT values based on the morning sounding do not AFH15-101 5 NOVEMBER 2019 205 necessarily imply severe weather will occur, but if SWEAT values remain high for the forecast sounding, then severe weather potential is high. 3.3.1.14. Showalter Stability Index (SSI). As discussed earlier in this chapter, the SSI is useful for determining lightning potential; this index works best in the Central United States, with well- developed systems. This index should only be used as a first indication of instability. It doesn‘t work well if a frontal surface or inversion is present between 850 mb and 500 mb. It also is not a good predictor of severe weather when low-level moisture is present below 850 mb. 3.3.1.15. Storm-Relative Directional Shear (SRDS). SRDS is used to measure the amount of directional shear in the lowest 3 km of the atmosphere; strong directional shear contributes significantly to storm rotation. 3.3.1.16. Storm Relative Helicity (SRH). Discussed in detail earlier in the chapter, SRH can indicate the likelihood of tornadogenesis; helicity has been found to correlate strongly with the development of rotating updrafts. Helicity is very sensitive to storm motion; storms that encounter boundaries or slow down can have radically different helicities than the general environment. SRH is not a tornado predictor by itself; it’s conditional on thunderstorm development. 0-3 km SRH values greater than 250 m2 s-2 suggest an increased threat of tornadoes with supercells, and 0-3 km SRH values exceeding 300 m2 s2 are usually found in environments that produce violent tornadoes. At lower levels, at least 100 m2 s-2 of 0-1 km SRH is needed for tornadogenesis. 3.3.1.17. Surface Cross Totals (SCT). Use SCT to predict severe potential for areas at high elevations. 3.3.1.18. Thompson Index (TI). Use TI to determine thunderstorm severity in mountainous regions, such as the Rockies. 3.3.1.19. Total Totals (TT). Use TT to forecast thunderstorm coverage and severity. This index is particularly good with cold air aloft. It may over- forecast severe weather when sufficient low-level moisture is not available. The TT index is the sum of the Vertical Totals and Cross Totals. 3.3.1.20. Vertical Totals (VT). Use VT in the western United States, the UK, and Western Europe to predict thunderstorm potential. 3.3.1.21. Wet Bulb Zero Height (WBZ). The WBZ is often a good indicator of hail and surface gusts greater than 50 knots when it lies between 5000 and 12,000 feet, and of tornadoes when it lies between 7000 and 9000 feet. It is not a good indicator in deep mT air masses, which naturally have high WBZs; hail or strong surface gusts rarely occur in these air masses outside the immediate vicinity of tornadoes. Multiple studies have indicated a strong correlation between the height of WBZ and the types of tornadoes that will occur; depending on the value, WBZ can help predict whether tornadoes will occur in isolated cases or form in groups. 206 AFH15-101 5 NOVEMBER 2019 3.3.2. Evaluation and Techniques. There are many data sources available to produce a forecast; atmospheric models and numerical analysis techniques, satellite, radar, conventional upper-air data, and a variety of software applications designed to help forecasters interpret these data. Deciding which tools and data to use in forecasting severe convective weather can be an overwhelming task; the following techniques can assist in the severe forecast process. Start with a knowledge of seasonal thunderstorm activity as described in regional climatology. 3.3.2.1. Synoptic evaluation for potential severe weather. Determine if the current and/or forecast weather pattern for the area of interest is favorable for severe convective weather pattern development. After initializing NWP model output, pay close attention to areas where favorable severe convective storm predictors stack with height. The more favorable conditions in a specific area, the greater the chance of development of severe thunderstorms. Use composite products to help stack significant features. When most predictors indicate strong potential for severe weather, seriously consider forecasting severe thunderstorms, tornadoes, strong winds, and/or hail. If predictors indicate weak potential, then consider forecasting non- severe thunderstorms. If indicators are mixed, consider forecasting non-severe thunderstorms with isolated or scattered severe thunderstorms. Finally, if low-level predictors are strong, weak upper-level diffluence is often sufficient to trigger severe weather, and if low-level predictors are marginal, strong upper-level diffluence is necessary to trigger severe convective storms. Incorporate local rules of thumb, forecast discussion bulletins, and various stability indices appropriate for the location into the decision-making process. Don’t base a forecast on a single tool when many are available! 3.3.2.2. Forecast products and techniques. Begin with the nearest representative Skew-T, and use the techniques described in this document to analyze the sounding for indications of convective instability in the air mass. Determine if the air mass is absolutely or conditionally unstable. Next, analyze the upper-air and surface products; upper-air analyses are not as useful for forecasting airmass thunderstorms as they are for forecasting severe thunderstorms, but they can often help. Local area work charts can play a key role in severe weather analysis, since they’re updated hourly; triggering mechanisms are often apparent on a shorter time scale. Table 3.7 provides a guide on key features on the standard-level charts, and Table 3.9 identifies key predictors to analyze, as well as their significance to severe weather occurrence. 3.3.2.3. Identifying tornado features. The first requirement for tornado prediction is a severe thunderstorm forecast; from there, a determination must be made whether tornadogenesis will occur. The magnitude of various parameters derived from the lowlevel wind and thermodynamic fields of the atmosphere are keys to tornado formation; the elements that contribute to tornadogenesis are strong storm- relative flow, strong vertical wind shear, strong low-level vorticity (i.e., strong low-level cyclonic circulation), potential for strong rotating updrafts and great instability or buoyancy. All of these elements are associated with supercells, which are known tornado producers. However, not all tornadoproducing thunderstorms are supercells. Several tools are available for determining whether conditions exist for tornadogenesis; these are shown in Table 3.5 with the parameters they measure, and what each tool is used to predict. The actual threshold values are listed in Table 3.2, Table 3.3, and Table 3.4 Several of these tools indicate storm type rather than tornado type or strength; knowing the expected storm type can indicate where AFH15-101 5 NOVEMBER 2019 207 tornadoes are likely to form within the storm, aiding severe storm METWATCH. Supercell tornadoes develop in the mesocyclone of classic and heavy precipitation supercells, and on the leading edge of the storm updraft in the vicinity of the wall cloud of low-precipitation supercells. Non-supercell tornadoes can occur in the flanking line of a supercell, during the growth stage in the updraft of “pulse” thunderstorms (strong, single-cell storms), along the gust front of multi-cell storms, and in strong updraft centers of multi-cell storms. Tornadoes in single cell and multi-cell storms are rare, and require exceptionally strong development to produce a tornado. 3.3.2.4. Identifying bow echo features. A bow echo is a line of storms that accelerates ahead of the main storm area; it forms from strong thunderstorms with a gust front. A strong downburst develops and the line echo wave pattern (LEWP) begins to “bow”. A well-developed bow echo or “spear head” is associated with the mature stage of the downburst; strong winds and tornadoes are possible near the bow. Figure 3.13 shows the evolution of the bow echo in a LEWP. As the downburst weakens, the line forms a comma shape, with a mesocyclone often developing on the north end of the comma, evident by a “hook” in the radar echo. At this point, tornadoes may still occur in the area of the mesocyclone, but the winds are decreasing. Strong to severe straight-line winds are likely to exist if four specific characteristics of the bow echo are present (See Figure 3.13 and Figure 3.14): the low-level echo configuration is concave downstream, weak echo channels exist, a strong reflectivity gradient along the leading edge of the concave-shaped echo exists, and the maximum echo top is over or ahead of the strong low-level reflectivity gradient. Figure 3.13. Line echo wave pattern (LEWP) and bow echo evolution. 208 AFH15-101 5 NOVEMBER 2019 Figure 3.14. Bow echo schematic and reflectivity example. AFH15-101 5 NOVEMBER 2019 Table 3.7. Product analysis matrix and reasoning. 209 210 AFH15-101 5 NOVEMBER 2019 Table 3.8. Identifying features of airmass thunderstorm development on upper air charts. AFH15-101 5 NOVEMBER 2019 211 3.3.2.5. Identifying microburst features. Microbursts are difficult to predict and detect, due to their small spatial scale (less than 4 km diameter), shallow vertical extent, and short life span. Ideal conditions for microburst formation occur when there is relatively warm and moist air in the low levels from the surface to the 700-600 mb (high theta-e), with relatively cool and dry air in the mid-levels from 600 mb to 400 mb (low theta-e). Microburst conditions also have relatively high CAPE values, (greater than 1000 J kg-1), although wet microbursts can occur with lower CAPE values as well. The following techniques can provide guidance on the potential for formation of microbursts. 3.3.2.5.1. Dry microburst forecast technique. Dry microbursts don’t only occur from thunderstorms; they may occur from any cumuliform cloud of appreciable vertical extent on top of a mixed (PBL) layer with a dry adiabatic lapse rate. This technique is a rapid way to determine dry microburst potential. 3.3.2.5.1.1. Using a morning sounding, look for a radiational temperature inversion at the surface with a depth of 40-50 mb. 3.3.2.5.1.2. Look for a dry adiabatic layer above this inversion that extends up to a level between 600 mb and 500 mb. 3.3.2.5.1.3. Find the average mixing ratio below the convective condensation level (CCL) and determine if it is less than 5 g/kg. 3.3.2.5.1.4. Look for relative humidity ≥ 70% above the dry adiabatic layer (usually at or above 600 mb). 3.3.2.5.1.5. Determine if the convective temperature (CT) for the day will be reached by mixing down the 850 mb temperature dry adiabatically to the surface, or using a model forecast sounding. If the CT will be reached, dry microbursts are possible. 3.3.2.5.1.6. Find the strongest wind speed within the dry adiabatic layer; if all of the above conditions are met, this wind speed may descend to the surface as part of a dry microburst. 3.3.2.5.2. The Dry Microburst Index (DMI). The DMI objectively quantifies dry microburst potential using mid-level lapse rates, as well as the presence or lack of midlevel moisture. The DMI is given by: DMI = Γ + (T – Td)700 − (T – Td)500. Where Γ is the 700-500 mb lapse rate in °C/km, and T and Td are the 700 mb and 500 mb temperatures and dew points, respectively, in °C. Environments in which the DMI is greater than 6 and instability exists aloft (CAPE greater than 50 J/kg) are considered favorable for dry microbursts. Large DMI values occur in environments with steep midlevel lapse rates, dry 750-700 mb air, and moist 550-500 mb air. Precipitation originating near 500 mb will fall into the dry air below, evaporating and enhancing downdraft magnitude, which may reach the surface as a microburst. Negative buoyancy is further increased if mid-level lapse rates are sufficiently steep. 3.3.2.5.3. Wind Index (WINDEX). The WINDEX is useful to forecast maximum potential wind gusts when both wet and dry microbursts are possible. It is a weighted formula to derive microburst potential based on melting (freezing) level height, lapse rates, and mixing ratios. See Table 3.9 for the equation for calculating the WINDEX. 212 AFH15-101 5 NOVEMBER 2019 Table 3.9. WINDEX Equation. 3.3.2.5.3.1. From a sounding, determine the height of the melting level (HM) in km AGL. 3.3.2.5.3.2. From the sounding, calculate the average mixing ratio (g/kg) in the lowest 1 km above the surface (QL). 3.3.2.5.3.3. Calculate the RQ value (divide the QL value from step 2 by 12). If this value is greater than 1, use 1. 3.3.2.5.3.4. From the sounding, calculate the lapse rate (Γ) in °C/km from the surface to the melting level. 3.3.2.5.3.5. From the sounding, find the mixing ratio (g/kg) at the melting level (QM). 3.3.2.5.3.6. Plug these values into the WINDEX equation to determine the maximum potential wind gust from a microburst if thunderstorms develop. 3.3.2.5.4. Wet microburst potential using VIL and Echo Tops values. This technique can provide up to 40 minutes of lead time predicting maximum downburst winds from pulse-type (single cell) thunderstorms. In order for this method to be effective, there must be a source of dry (dew-point depression greater than 18°C), potentially cold air between 400 and 500 mb, along with reflectivity values 55 dBZ or greater (allowing for sufficient moisture for entrainment of the air parcel to produce negative buoyancy through evaporative cooling. To predict wind gust potential with this method, find the Echo Top and VIL values of the storm cell and refer to Table 3.10 for the maximum downburst wind potential in knots. This technique won’t work when VIL values are corrupted due to hail contamination, or when thunderstorms are too close to the radar, causing Echo Top estimates to be erroneously low. This technique also works poorly for thunderstorms over 125 NM away from the radar. The VIL and Echo Top method only works for pulse-type air-mass thunderstorms; it won’t provide accurate estimates for multi-cell or supercell storms. AFH15-101 5 NOVEMBER 2019 213 Table 3.10. Wet microburst potential table, using VIL and Echo Tops. 3.3.2.5.5. Wet microburst potential using the Atkins and Wakimoto (1991) method. Find the maximum equivalent potential temperature (theta-e) value (in degrees Kelvin) at the surface, then find the minimum equivalent potential temperature (theta-e) value (in degrees Kelvin) in the mid-levels, between 600 and 400 mb. Calculate the difference between the values; if the difference is greater than 20K, wet microbursts are likely. If the difference is less than 13K, wet microbursts are unlikely. 3.3.2.5.6. Wet microburst potential using the Microburst-Day Potential Index (MDPI). This method is not a tool for thunderstorm forecasting; it assumes that thunderstorms will develop on the day in question. Find the maximum theta-e value between the surface and 850 mb (in degrees Kelvin), then find the minimum theta-e value between 660 mb and 500 mb (in degrees Kelvin). Calculate the difference between the values and divide by 30; if the MDPI is greater than 1, wet microbursts are likely. If the value is less than 1, wet microbursts are unlikely. 3.3.2.5.7. Wet microburst potential using the Wet Microburst Severity Index (WMSI). This index was developed by case studies of 35 wet microbursts in the eastern and central United States; it may need adjustment for other regions of the world. Find the maximum theta-e value at the surface (in degrees Kelvin), then find the minimum thetae value (in degrees Kelvin) in the middle levels between 600 mb and 400 mb. Calculate the difference between these two values. Find a representative value of CAPE in the forecast area, multiply the CAPE value by the theta-e difference, then divide the result by 1000; this is the WMSI value. WMSI values between 10 and 49 indicate wind gusts less than 49 knots, WMSI values between 50 and 79 indicate wind gusts from 35-49 knots, and WMSI values greater than 80 equate to wind gusts greater than 50 knots. 3.3.2.6. Boundaries and boundary interaction features. 214 AFH15-101 5 NOVEMBER 2019 3.3.2.6.1. Satellite. As diurnal heating occurs, cumulus clouds will often form into cloud streets (over land) oriented with the gradient wind flow. Look for clear areas forming in the flow; these identify sea-breeze fronts, lake breezes, and outflow boundaries. The leading edge of these boundaries between clear areas and cloud streets is highly favorable for development. Similarly, the boundary between cloud-free areas and fog or stratus broken/overcast areas are prime for development as clouds burn off. When outflow boundaries intersect, convection is almost guaranteed if the air mass is unstable or conditionally unstable. 3.3.2.6.2. Radar (Figure 3.15). Sea breeze boundaries and other discontinuities in lowlevel flow can usually be identified on base reflectivity displays. The sea breeze will appear as a thin line of low intensity returns, parallel to coastlines. Convection is likely to form on these lines when the convective temperature is reached. Figure 3.15. Sea breeze on a base reflectivity product. 3.3.2.6.3. Streamline analysis and sea breeze onset. Streamline analyses can be useful when combined with current satellite and radar analyses; create a composite product to identify locations of streamline-confluent asymptotes, sea/lake breezes, and outflow boundaries. Mark past locations of these boundaries, and determine speed and direction of boundary movement to project when and where these boundaries will intersect. The intersections are almost certain to result in air-mass thunderstorms. If thunderstorms are present along the boundaries already, severe weather (usually severe wind gusts) is possible. Tornadoes and hail are unlikely, unless strong upper-level support is evident. 3.3.2.7. Severe thunderstorm checklist. No two thunderstorm situations are alike; there are varying degrees of intensity for each parameter, and the combinations of parameters produce individual storm events. A foolproof, all-inclusive checklist for severe weather is impossible to create, so the steps below are an outline of the forecast reasoning process. Incorporate local rules of thumb and stability thresholds to fine-tune these steps for your individual location: AFH15-101 5 NOVEMBER 2019 215 3.3.2.7.1. Identify the current weather regime and expected synoptic type (dryline, frontal, overrunning, cold core, squall line, or airmass thunderstorm). 3.3.2.7.2. Analyze available model data, and tailor the analysis. 3.3.2.7.3. Are elements for severe weather present? Refer to Table 3.2, Table 3.3, Table 3.4, Table 3.5, Table 3.7, and Table 3.8 for features associated with severe weather elements. 3.3.2.7.4. Analyze current and forecast Skew-Ts and calculate stability indices appropriate for the weather pattern and station. Do they indicate severe weather potential? 3.3.2.7.5. Examine current and forecast hodographs; what types of storms can be expected from the wind shear pattern? 3.3.2.7.6. Based on the above steps, what type of severe weather can be expected, if any: winds, hail, or tornadoes? 3.3.2.8. Forecasting convective wind gusts. This section presents a variety of methods to forecast convective wind gusts; each technique is designed to forecast winds under different conditions. Each method requires a current or forecast Skew-T. 3.3.2.8.1. T1 gust computation – method 1. The T1 method is designed for scattered thunderstorms in the vicinity of the forecast location; method 1 applies when there’s an inversion present within 150-200 mb of the surface and is not susceptible to breaking from surface heating. 3.3.2.8.1.1. Project the moist adiabat from the warmest point of the inversion to the 600 mb level. 3.3.2.8.1.2. Calculate the temperature difference (in °C) between the moist adiabat and the dry-bulb temperature trace at 600 mb. Label this point as T1. 3.3.2.8.1.3. Refer to Table 3.11; the value for T1 is considered to be the average gust speed. Add 1/3 of the lowest 5000 foot mean wind speed to the table value to find the maximum gust speed. Wind gust direction can be determined from the mean wind direction in the layer between 10,000 and 14,000 feet above the local terrain. 3.3.2.8.2. T1 gust computation – method 2. This method applies when there’s no inversion present, or the inversion is relatively high (more than 200 mb above the surface). 3.3.2.8.2.1. Forecast the maximum surface temperature. 3.3.2.8.2.2. Project the moist adiabat from the maximum temperature to the 600 mb level. 3.3.2.8.2.3. Calculate the difference between the moist adiabat and the dry-bulb temperature trace at 600 mb. Label this point as T1. 216 AFH15-101 5 NOVEMBER 2019 3.3.2.8.2.4. Refer to Table 3.11; the value for T1 is considered to be the average gust speed. Add 1/3 of the lowest 5000 foot mean wind speed to the table value to find the maximum gust speed. Wind gust direction can be determined from the mean wind direction in the layer between 10,000 and 14,000 feet above the local terrain. Table 3.11. T1 convective gust potential. 3.3.2.8.3. T2 gust computation. The T2 method is designed to determine gust potential for intense squall lines or numerous thunderstorms. 3.3.2.8.3.1. Find the WBZ height – the point where the wet bulb curve crosses the 0°C isotherm. 3.3.2.8.3.2. Project the moist adiabat through the WBZ to the surface, and note the surface temperature in °C at that point. 3.3.2.8.3.3. Subtract the moist adiabat temperature from the surface dry bulb (or projected maximum) temperature; this is the T2 value. 3.3.2.8.3.4. Refer to Figure 3.16; follow the T2 value up to the intersection of the three curves. The first intersection point is the minimum gust, the middle intersection point is the average gust, and the upper intersection point is the maximum gust. Wind gust direction can be estimated from the mean wind direction between 10,000 and 14,000 feet. 3.3.2.8.4. Snyder method gust computation. The Snyder method is best used to forecast average wind gusts with airmass or “pulse” type thunderstorms. 3.3.2.8.4.1. Find the WBZ height. 3.3.2.8.4.2. Forecast maximum temperature at time of thunderstorm occurrence (in °F). 3.3.2.8.4.3. Lower the WBZ to the surface along the moist adiabat, and find the surface temperature (in °F) at the intersection point (the “down rush” temperature). 3.3.2.8.4.4. Subtract the step 3 value from the step 2 value (down rush temperature minus forecast max temperature). AFH15-101 5 NOVEMBER 2019 217 3.3.2.8.4.5. Find half the average wind speed in the 10,000 foot layer centered on the WBZ height (5000 feet below the WBZ height to 5000 feet above the WBZ height). 3.3.2.8.4.6. Add the step 4 value to the step 5 value; this is the average gust potential for air-mass thunderstorms in the current environment. Figure 3.16. T2 gust computation chart. 3.3.2.9. Hail forecasting. Hail is a common, microscale phenomenon associated with strong thunderstorms; the key factors to determine are whether or not the hail in a thunderstorm will reach the surface, and if so, what size hailstones will fall to the ground. 3.3.2.9.1. Forecasting hail occurrence. The first step in hail forecasting is to determine whether or not it will occur; this technique is an objective method derived from numerous case studies of severe thunderstorms over the Midwestern United States. The method determines the cloud depth ratio, and then correlates the ratio and freezing level to occurrence or non-occurrence of hail. 3.3.2.9.1.1. From a Skew-T diagram, calculate the Convective Condensation Level (CCL), Equilibrium Level (EL), and Freezing Level (FL). 3.3.2.9.1.2. Determine the cloud depth ratio using the following equation: (CCL – FL)/(CCL – EL) 218 AFH15-101 5 NOVEMBER 2019 3.3.2.9.1.3. Plot the cloud depth ratio and freezing level on Figure 3.17. If the plot is below the line, expect hail; if it’s above the line, hail is not likely to occur. Figure 3.17. Hail prediction chart, using cloud depth ratio and freezing level. 3.3.2.9.2. Forecasting hail size. If hail is likely to occur, this technique can assist in determining hail size potential. 3.3.2.9.2.1. From a Skew-T diagram, calculate the CCL, then follow the saturation mixing ratio line to its intersection with the temperature trace (Figure 3.18, point A). 3.3.2.9.2.2. The intersection of the -5°C isotherm and the sounding is point BH; note the pressure level where this intersection occurs. 3.3.2.9.2.3. From point A, follow the moist adiabat up to the point BH pressure level; this is point B’ (Figure 3.18). 3.3.2.9.2.4. Calculate the temperature difference (in °C) between points BH and B’. 3.3.2.9.2.5. From point BH, follow the dry adiabat to the CCL; this is point H’. Calculate the temperature difference between BH and H’. 3.3.2.9.2.6. Refer to Figure 3.19 to forecast preliminary hail size; find the BH-B’ temperature difference on the x-axis, the BH-H’ temperature difference on the yaxis, and plot the intersection – this will provide an initial estimate of hail size potential. 3.3.2.9.2.7. Find the WBZ height; if it’s less than 10,500 feet, the preliminary hail size computed in step 6 will be the final hail size. If the WBZ height is above 10,500 feet, refer to Figure 3.20 and use the preliminary hail size and WBZ height to determine final hail size potential. AFH15-101 5 NOVEMBER 2019 Figure 3.18. Hail size parameters on the Skew-T diagram. Figure 3.19. Preliminary hail size nomogram. 219 220 AFH15-101 5 NOVEMBER 2019 Figure 3.20. Final hail size nomogram – only use if WBZ height is above 10,500 feet. 3.3.2.9.3. Forecasting hail size using VIL density. VIL is a function of radar reflectivity data converted into an equivalent liquid water content value that is largely based on the sensed hydrometeor/target (rain drops, hail, snowflakes, etc.) size distribution. VIL is also a function of a “reflectivity factor” that has target diameter as its primary component (number of targets is a much smaller component). As target diameter increases, so does the reflectivity value. In fact, reflectivity increases exponentially as target diameter increases, so a small increase in target size results in a much more significant increase in the reflectivity value obtained. VIL values increase exponentially with increasing reflectivity, so large VIL values mean high reflectivity, which in turn often indicates the presence of large targets (i.e., hail) within the volume scan. However, VIL alone is not an adequate hail indicator because it is air mass dependent, neglects storm depth, and may be inaccurate for storms very close to or far away from the radar. VIL Density “normalizes” the VIL values using the height of the thunderstorm (i.e., echo top) as indicated by radar. This eliminates the air mass dependency of stand-alone VIL values, and is useful to identify high reflectivity thunderstorms relative to their height. This is important because hail growth is often a function of cloud height/depth. Table 3.12 shows approximate relationships between VIL density and observed hail sizes; use the VIL density values from observed storms on radar to obtain a quick, real-time hail estimate. VIL Density has the following limitations: 3.3.2.9.3.1. VIL Density only indicates hail aloft. Hailstones that reach the surface may be much smaller than that indicated by the VIL Density values. Examining WBZ levels in tandem with VIL Density values is useful to mitigate this limitation. High VIL Density values and a WBZ Level between 5,000-11,000 feet indicate a higher potential of large hailstones reaching the surface. AFH15-101 5 NOVEMBER 2019 221 3.3.2.9.3.2. VIL and VIL Density values will be quite accurate for slow moving thunderstorms with limited horizontal tilt such as found during the summer months in the mid-latitudes, but less accurate for fast-moving, tilted thunderstorms in environments of strong VWS. This may result in VIL values being averaged and lower than they actually are, resulting in smaller VIL Density values than reality. 3.3.2.9.3.3. Echo tops associated with reflectivity values greater than 18 dBZ may not be accurate because of scanning strategies chosen. In general, Volume Coverage Pattern (VCP) 11 does a better job of estimating VIL and echo top values than VCP 21. Table 3.12. VIL density and hail size. 3.3.2.10. Onset of typical thunderstorms. Predict thunderstorm onset at the time when the convective temperature is forecast, or maximum solar heating is expected. Predict formation along confluent streamline asymptotes and discontinuities in the flow such as sea breezes, outflow boundaries, and lake breezes. 3.3.2.11. Onset of severe thunderstorms. Hail, severe winds, and tornadoes are less common with air mass thunderstorms – for severe weather to occur, at least one of the following must be present: 3.3.2.11.1. Cold and/or dry air aloft. 3.3.2.11.2. Short-wave trough at 500 mb. 3.3.2.11.3. Positive vorticity advection (PVA). MARK D. KELLY, Lt Gen, USAF Deputy Chief of Staff, Operations