Fire behavior is controlled by three interacting components: fuels, weather, and topography. Fuels provide the energy source for fire. Fuel availability, which depends on both fuel arrangement and fuel moisture, determines if fires will burn as ground, surface, or crown fires. Weather elements, such as temperature, relative humidity, wind, precipitation, and atmospheric stability, also combine to influence fire behavior by regulating fuel moisture and rate of spread. Topography can influence fire indirectly, by mediating wind patterns, or directly- fires burning upslope spread faster than fire burning on flat land.

The variety of fuel, weather, and topographical conditions that exist in the South create fires that vary in the amount of fuels that burn, the rate at which these fuels burn, the depth of burns, and whether living plants become fuel. This variation in fire behavior, in turn, influences the effects of fire on natural communities and people. This section of the encyclopedia provides a background on how fuels, weather, and topography influence fire behavior.

• Fire, Weather, and Fire Behavior- Overview

• Fuels: This section defines fuel categories (ground, surface, ladder, and canopy fuels); explains how physical fuel properties, such as size, shape, loading, and arrangement, combine to influence combustion and fire behavior; explains how chemical fuel properties affect heat content and flammability; describes the importance of fuel moisture and other factors that control fuel availability.

• Fire Weather: This section defines important components of fire weather and explains how each of these components influence fire behavior and risk of wildfire.

• Fire Behavior: This section explains how fuels, weather, and topography combine to influence fire behavior. It includes a discussion of combustion and heat transfer, fire intensity, fire growth and spread, extreme fire behavior, and fire behavior prediction systems.

It is important to understand what controls fire behavior and how to predict it. This knowledge will help predict fire effects, conduct prescribed burns, predict wildfire risk, and control wildfires. Fire behavior is controlled by three interacting components: fuels, weather, and topography. Fuels provide the energy source for fire. Fuel availability which depends on both fuel arrangement and fuel moisture determines if fires will burn as ground, surface, or crown fires. Weather elements such as temperature, relative humidity, wind, precipitation, and atmospheric stability combine to influence fire behavior by regulating fuel moisture and rate of spread. Topography can influence fire indirectly, by mediating wind patterns, or directly- fires burning upslope spread faster than fire burning on flat land.

It is important to understand the physical-chemical process of fire to understand how heat is generated by fire. Fire releases heat through combustion. Oxygen, heat, and fuel- often called the fire triangle- must be present in the proper ratio for a fire to ignite and sustain combustion. Once a fire has ignited, the heat must be transferred to surrounding fuel in order for the fire to grow and spread. This occurs through one of several heat transfer processes, usually convection, radiation, and/or conduction, although vaporization and mass transport may also play roles.

Once a fire has ignited, its shape and rate of spread will continually change. Rarely, a fire can continue to increase its rate-of-spread and intensity, resulting in extreme fire behavior- a level of fire behavior that goes beyond human methods of fire control and prediction.

Characterizing flame attributes such as flame height, length, depth, angle, and char height can help predict fire effects and make comparisons among different fires possible. Fire intensity, which describes the rate of heat release, and rate of spread play significant roles in characterizing fire behavior.

Computer models can be used to predict fire behavior based on differences in fuels, weather, and topography. These fire behavior prediction systems are used to support fire management decisions, as a training tool to improve fire management skills, and can help display and explain fire behavior and fire management strategies to the general public. Fire danger rating systems produce qualitative and/or numeric indices of fire potential based on fuels, topography and weather. These rating systems allow fire managers to estimate present and future fire danger for a given area. Both fire behavior prediction systems and fire danger rating systems require mathematical descriptions of fuel models and their respective fuel properties as input.

Fuel is all living and dead plant material that can be ignited by a fire. Fuel characteristics strongly influence fire behavior and the resulting fire effects on ecosystems. Fires vary widely in the kind of fuels that burn (e.g., live vs. dead fuels, surface vs. ground fuels), the total amount of fuels that burn, and the rate or intensity at which these fuels burn. These characteristics of fuel consumption, in turn, determine peak temperatures reached, the duration of heat, and the stratification of heat above and below the soil surface (Miller 1994). The following sections discuss concepts that will help users understand how fuels affect fire behavior.

• Fuel Categories defines ground, surface, ladder, and canopy fuels and explains factors controlling the availability of each fuel type.
• Physical Fuel Properties explains how fuel size, shape, loading, and arrangement combine to influence combustion and fire behavior. Several techniques for measuring fuel loads are also described.
• Chemical Fuel Properties describes how the chemical composition of fuels affects their heat content and flammability.
• Fuel Moistureexplains how fuel moisture determines fuel availability, what controls fuel moisture, defines time-lag classes, and describes several techniques for measuring fuel moisture.
• Fuel Availability and Consumption explains what factors control the availability of both live and dead fuels.
• Plant Flammability discusses species-specific factors that control the flammability of live plants and presents flammability indices for common southern plant species.

Fuels can be classified into four broad categories based on their vertical distribution:

• Ground fuels: organic soils, forest floor duff, stumps, dead roots, and buried fuels
• Surface fuels: litter layer, downed woody materials, and dead and live plants to 2 m (6 ft.) height
• Ladder fuels: vine or liana fuels and draped foliage fuels
• Canopy fuels: tree crowns

These fuel categories are not to be confused with the fuel types used in fuel models (such as grasses, brush, timber litter, and logging slash). Fuel models are more specific classes of fuels used in fire behavior modeling.

Ground fuels are those forest fuels that lie below the litter layer or within the soil, including organic soils, forest floor duff, stumps and dead roots, and buried fuels. Ground fuels can ignite and smolder for days to months following flaming front passage. Ground fires produce persistent and harmful smoke and can re-ignite surface fuels making them a bane for fire managers.

The forest floor is the layer of organic matter overlying the mineral soil and has both surface and ground fuel components. The forest floor fuel complex contains distinct horizons, each with different moisture relationships, particle sizes, chemical composition, densities, and depths. The surface fuel component of the forest floor is the litter (Oi) horizon. The ground fuel component, duff, is beneath the litter horizon. It is comprised of the fermentation (Oe) and humic (Oa) horizons. In long-fire interval ecosystems the duff layer can become well-developed, however in frequently burned systems it may be intermittent or nonexistent. Duff is created by litter decomposition, so many volatile compounds are lost, particle sizes are reduced, and it is shaded by the overlying litter horizon. Similar to 1,000-hour timelag fuels, duff is slow to absorb moisture. Therefore, when duff moisture is low, smoldering phase combustion often consumes this horizon, resulting in high fire severity and copious amounts of smoke.

Organic soils are important forest fuels in several southeastern ecosystems. Organic soils contain the duff layer overlying a variety of soils (see earlier discussion) and true histosol organic soils. Histosols are dark-colored soils consisting of large amounts of organic peat and muck, underlying poorly-drained forested and nonforested wetlands (e.g., cypress domes, pitcher plant bogs, and bay swamps). Available fuel in organic soils is defined by three factors: moisture, packing, and mineral soil content (Frandsen 1987). Increases in any of these factors decreases flammability and retards combustion. However, following extended droughts, organic soils can ignite and burn for days to months, often smoldering beneath the surface (so called “muck fires”). Organic soil fires are serious concerns in many southeastern wetland communities; they are difficult to control, and have serious ecosystem effects (see: Prescribed Burning in Organic Soils).

See also: Moisture Content of Ground Fuels.

Surface fuels are the primary fuel of interest for fire behavior in most southeastern ecosystems (Wade et al. 2000). Surface fuels include understory plants < 2 m (6 ft.) tall (dead and alive), the litter layer, downed woody materials, and often midstory tree and shrub fuels. Surface fuel availability for consumption is determined by moisture content, particle size, horizontal continuity, compactness, and fuel type (particularly fuels with high volatile compounds). Under most burning conditions in most southeastern ecosystems these fuels carry surface fires.

The understory is the layer of living and dead vegetation from the soil surface to 2 m (6 ft.) tall. Many southeastern ecosystems (e.g., open pine savannas and forests, freshwater marshes, pitcher plant bogs, prairies) contain a dominant understory with abundant grass, forb, small woody shrub and litter fuels. Both grasses and their allies (sedges and rushes) and forbs have high surface area-to-volume ratio, low fuel moisture, are within the flaming zone of most surface fires, and retain abundant dead leaves making them ignite and combust rapidly (exceptions to this are succulents and large-leaved species). Understory fuel availability in southeastern ecosystems is controlled by fuel moisture, horizontal fuel continuity, and fuel loading.

Small woody shrubs can be important understory surface fuels (Blackmarr and Flanner 1975, Hough and Albini 1978). Pocosins, flatwoods, sand pine scrubs, and bogs contain large loadings of shrubby fuels. Many southeastern shrubs have high surface area-to-volume (e.g., saw-palmetto, Serenoa repens), high volatile contents (e.g., gallberry, Ilex glabra), grow within the flaming zone of surface fires, and are highly flammable. In some ecosystems, shrubs and small trees grow into the midstory (between 2 and 5 m; 6 and 16 ft.) and carry surface fires into lower canopy fuels. Midstory fuel availability is regulated by vertical fuel continuity, fuel moisture, and fire behavior. Low-intensity fires with low flame lengths often don’t ignite midstory shrub fuels.

The forest floor is the layer of organic matter overlying the mineral soil and has both surface and ground fuel components. The surface fuel component of the forest floor is the litter (Oi) horizon. The ground fuel component of the forest floor is the duff layer. Litter horizons are fuels in almost all forested southeastern ecosystems, and are therefore somewhat diverse in their composition and structure. Most litter horizons contain recently deposited litter, small woody fuels (10-, 100-, and few 1,000 hour timelag fuels), cones, and other dead plant parts. Litter fuels have reduced volatile content, low fuel moisture content (often 5 to 15%), and are usually loosely packed. Surface fires can be carried solely by litter fuels. Litter fuels may also ignite live understory fuels, pre-heat larger woody fuels, and initiate smoldering of underlying ground fuels, if present. Forest floor fuel availability is determined primarily by fuel moisture content and fuelbed bulk density. Separation of available and unavailable fuel is made on depth to moisture, with all dry fuel included as surface fuel and the remaining wet included as ground fuel.

Understory and shrub fuels are measured using quadrat, point-quarter center, and line transect sampling methods (see Measuring Fuel Loads). Loads (measured in dry kg/m² or lb/acre) are calculated and extrapolated to larger areas or can be input into fire behavior models (e.g., BEHAVE). Forest floor surface fuels are measured by harvesting small quadrats (in kg/m² or lb/acre, and drying for moisture content) and by determining fuelbed bulk density (in kg/cm² or lb/ ft³).

Ladder fuels are those that provide vertical continuity between understory or midstory surface fuels and canopy fuels. Ladder fuels consist of vine or liana fuels, draped foliage fuels, and hanging broken branches. Most surface fires in southeastern ecosystems involve isolated ladder fuels, though in particular circumstances ladder fuels can accumulate and lead to high severity fires.

Vine fuels include several southeastern species that are important in surface fires, such as yellow jessamine (Gelsemium sempervirens), greenbriers (Smilax spp.) and wild grape (Vitus spp.). Vines ascend trees and shrubs creating vertically continuous fuels. Dead and live foliage, stems, and flower structures have low fuel moisture, are bathed in convective heat, and contain volatile compoundscrown fires. Vine fuel availability is governed by fuel moisture, flame height, windspeed, and the live-to-dead ratio. Vine fuels are usually measured as presence/absence, height in canopy, vine loads, and live to dead ratios.

Draped dead foliage (especially pine needle litter) on vines and living or dead shrubs is another important ladder fuel type. Draped fuels have very low fuel moisture (wind, sun, and humidity effects are increased) and are highly flammable. Draped fuels increase the height of the combustion zone, linking understory and midstory fuels to canopy fuels. Southeastern pine plantations and long-unburned forested ecosystems with well-developed vine and/or shrub layers are especially prone to draping.

Hanging broken branches become important ladder fuels in forests following hurricanes, tornadoes, ice storms, and other disturbances.

Canopy fuels are the crowns of trees that form the overstory. The receptivity of the canopy fuels to crown fire is based primarily on three factors: canopy base height, canopy bulk density, and, to a lesser degree, foliar moisture content (Fieldhouse and Dickinson 2003). Canopy base height relates the bottom of the overstory tree crowns to the top of the understory fuel bed and ladder fuels. Canopy bulk density is a measure of the amount of fuel contained in a unit volume of the canopy. High bulk densities present large fuel loads for a fire.

Canopy or crown fuels are typically not consumed during fires in the southeastern US except in isolated cases of "torching" which affect individual trees. Crown foliage is commonly scorched, but rarely is it consumed (i.e., combusted) in crown fires. Particular exceptions are the stand-replacing fires common in sand pine scrub forests in central Florida, in stand-replacing fires in non-indigenous melaleuca forests in south Florida, and limitedly in Table mountain pine forests in the southern Appalachians. General exceptions to this statement occur in fires, either prescribed or wildfire, with extreme fire behavior (caused by low moisture levels, erratic winds, or high fuel loadings).

The primary physical fuel properties influencing combustion and fire behavior are size, shape, loading, and arrangement. Here, we define these four properties and discuss the physical properties of typical fuels found in the southeastern US.

Size

Particle size is one of the most important fuel characteristics affecting combustion and fire behavior (Byram 1959). Large particles have high heat capacities, requiring more heat to ignite and combust the particle. Smaller particles have low heat capacities, so they require smaller amounts of heat energy for ignition and combustion. For dead fuels, particle size is also related to the rate at which fuel moisture content changes, and therefore size classes of fuels are also referred to as timelag classes. Different time-lag classes burn differently: 1-hour fuels (needle litter, hardwood leaves) ignite quickly and combust at rapid rates. Progressively larger particles (10-, 100-, 1000-hour and larger fuels) require more heat for ignition and combustion. Fires usually start and spread in dead fines fuels (< ¼ in. diameter), which ignite increasingly larger size classes of fuels. If fine fuels are reduced or missing, a fire may not ignite or spread.

Shape

Fuel shape (surface area-to-volume ratio) is related to particle size: the more finely divided the material, the higher the ratio. Fuel surface area is measured in cm²/ m³ (or ft²/ ft³). Fuels with high surface area-to-volume ratios (pine needle litter, most foliage fuels) have lower heat capacity and require less pre-heating for ignition (Byram 1959). The increased surface area of these fuels provides more surface area for heat oxidation and combustion. Further, these fine fuels dry out and ignite more rapidly than coarser fuels. Small surface area-to-volume ratio fuels (downed logs and other 1,000-hour fuels) resist ignition and combust slowly. There are many examples of fuels from southeastern ecosystems that have high surface area-to-volume ratios.

Loading

Fuel loading is the amount of live and dead fuel, expressed in weight per unit area (kg/m² or tons/acre). Total fuel is all fuel, both living and dead, present on a site. Available fuel is the amount of fuel that will burn under a specific set of fire conditions (Pyne et al. 1996). Fuel loadings are usually grouped by particle size class (or timelag classes).

Fuel loading is an important characteristic of southeastern fuel complexes. Fuel loads vary considerably depending on site productivity, recent disturbance history, and fire regime. While the generally warm and humid Southeastern climate provides optimal growing conditions, some systems are more productive than others. Fuel production varies from low in the xeric sand pine scrubs, Appalachian ridgeline ecosystems, and Piedmont granite outcrops, to high in mesic pine forests and most wetland communities. Recent disturbances may increase or decrease existing fuel loads, by either removing fuels (in the case of fire) or adding new fuels in the form of coarse woody debris (in the case of hurricanes). For this reason, long-unburned stands typically have higher fuel loads than stands managed with frequent prescribed fire. In these long-unburned stands, midstory and overstory fuels generally increase at the expense of fine fuels in the understory.

See Methods for Measuring Fuel Loads.

Arrangement

Fuel arrangement is another important physical property of fuels. Both the packing ratio and fuel placement describe different aspects of fuel arrangement:

Packing ratio is a measure of the compactness of the fuel bed. It is expressed as a percentage of the fuel bed composed of fuel, with the remainder being air space. Densely packed fuels prevent moisture evaporation and oxygen diffusion into the fuelbed, thereby suppressing ignition and flaming combustion. Conversely, loosely packed fuels allow rapid evaporation and oxygen diffusion, and hence rapid ignition and flaming combustion. However, fuels that are very open can burn slowly because little heat is transferred among widely spaced particles. For every size of fuel particle, there is an optimum packing ratio at which heat transfer and oxygen produce the most efficient combustion (Burgan and Rothermel 1984). Draped pine needles and upper layer forest floor litter are examples of fuels with low packing ratios. Live branchwood is a classic example of densely packed fuel. Fuel bulk density is a related measure of the compactness of a fuel or fuel bed. It is calculated by dividing the weight per unit area by the fuel bed depth, and is expressed as g/cm³ or lb/ft³. In general, the higher the bulk density of fuel is, the higher the spread rate (Miller 1994).

Fuel placement and fuel continuity describe the horizontal and vertical distribution of fuels (Pyne et al. 1996). Fuels placed within the flaming zone are available for combustion, whereas fuels out of the combustion zone are not. Fuels with horizontal and/or vertical continuity pre-heat adjacent fuels. Conversely, fuels lacking continuity do not transmit heat to adjacent fuels. Horizontal continuity is a critical factor for surface fires; bare patches and patches of sparse vegetation act as fuelbreaks. In crown fires, vertically continuous fuels facilitate crown ignition and crown-to-crown horizontal continuity sustains crown fire (Pyne et al. 1996).

Many fuels typical of southeastern ecosystems are highly flammable due to their high surface area-to-volume ratios (Brown 1970). In particular, long-needle pine litter and grasses have very high surface areas and are responsible for carrying surface fires in many southern ecosystems. This page provides examples of plant species with high surface area-to-volume ratios.

Grasses, Sedges, Rushes

Grasses and their grass-like allies (sedges and rushes) have extremely high surface area-to-volumes and are highly flammable fuels that carry surface fires in many southeastern ecosystems. Notable examples of these flammable fuels are:

• wiregrasses (Aristida stricta and others),

• bluestems and broom sedges (Andropogon and Schizachyrium spp.),

• giant cane (Arundinaria gigantea),

• panic grasses (Panicum and Dichanthelium spp.),

• muhly grasses (Muhlenbergia spp.),

• the invasive cogon grass (Imperata cylindrica), and

• sedges (particularly sawgrass, Cladium jamaicense) and rushes.

Conifers

Conifers in the southeastern US contain a large group of species with highly flammable litter (Fonda 2001) due to the high surface area-to-volume of conifer needles.

• Long-needle pines have the highest surface area-to-volume ratios of conifers. Examples are: longleaf pine (P. palustris), south Florida slash pine (P. elliottii var densa), slash pine (P. elliottii var elliottii), loblolly pine (P. taeda), pond pine (P. serotina).

• Short-needle pines have lower surface area-to-volume ratios but are still flammable. Examples include: shortleaf pine (P. echinata), Ocala sand pine (P. clausa var. clausa), Choctawhatchee sand pine (P. clausa var. immuginata), spruce pine (P. glabra), Virginia pine (P. virginiana), Table Mountain pine (P. pungens), eastern white pine (P. strobus), and pitch pine (P. rigida).

• Other southeastern conifers have reduced surface area-to-volume, but still much higher than small woody fuels (10- and 100-hour timelag fuels) and many hardwood foliage fuels. Examples include: eastern redcedar (Juniperus virginiana), southern redcedar (J. silicicola), Atlantic white-cedar (Chamaecyparis thyiodes), bald cypress (Taxodium distichum), pond cypress (T. ascendens), eastern hemlock (Tsuga canadensis), Florida torreya (Torreya taxifolia), Fraser fir (Abies fraseri), and red spruce (Picea rubens).

Forbs, shrubs, palms, vines, ferns, lichen, and bromeliads

Other southeastern fuels that are notable for their surface area-to-volume and resulting flammability are forbs, palms, lichens, ferns, vines and bromeliads.

• There is a great diversity of forbs and shrubs with high surface area-to-volume ratios and make up a large portion of the available fuels in several southeastern ecosystems.

• Palm leaves have high surface area-to-volume and are highly flammable; saw-palmetto (Serenoa repens), cabbage palm (Sabal palmetto), and scrub palm (Sabal etonia) are notable examples.

• Lichen fuels (foliose lichens; the reindeer “mosses” Cladonia and Cladina spp. and draped Usnea spp.) have extremely high surface area-to-volume, and combust readily where they dominate patches in pine sandhill and sand pine scrub ecosystems (Brown 1970).

• Understory ferns (particularly bracken fern, Pteridium aquilinum) also have high surface area: volume and combust readily in southeastern surface fires.

• Several draped vines (many Smilax spp. and yellow jessamine, Gelsemium sempervirens) have high surface area-to-volume and contribute to surface fires and localized crown torching.

• Spanish moss (Tillandsia usneoides) and Tillandsia recurvata can be important firebrand fuels, contributing to long-distance spotting (Clements 1976) and spill-over.

Hardwoods

Hardwood foliage has a great diversity of surface area-to-volume, so flammability varies drastically in these fuels. Generally speaking, hardwood litter can burn under dry conditions, but is much less flammable than needle litter and standing dead herbaceous fuel.

• Among hardwoods, most oaks [notably turkey oak (Quercus laevis), post oak (Q. stellata and Q. margaretta), and white oak (Quercus alba)], many other deciduous tree and shrub species [black cherry (Prunus serotina), blackgum (Nyssa sylvatica), hawthorns (Crataegus spp.), and heaths) have high surface area-to-volume and are the most flammable.

• Other deciduous hardwoods and most evergreen hardwood species [e.g., live oaks (Quercus virginiana and Q. geminata) and southern magnolia (Magnolia grandiflora)] have lower surface area-to-volume and require more pre-heating for ignition and combustion.

There are other species-specific characteristics that determine flammability in addition to surface area-to-volume ratios, such as the concentration of organic volatiles. For information on the flammability of individual plant species that take into consideration all of these characteristics, see Plant Flammability.

Along with fuel moisture, fuel loading is among the most important variables affecting fire behavior and consumption (Hough 1968). Fuel loading is a critical input in all fuel models, with bearing on flaming and smoldering fire combustion and fire severity. Estimating fuel loads is therefore a critical step in planning prescribed fires or assessing risk of fire danger. Fuel loading is a function of site productivity, decomposition rate, and time since last disturbance (e.g., fire, harvesting, past land-uses). Along with the rate of fuel accumulation, the standing crop of fuel (fuel loading) varies over the southeastern landscape.

Field sampling methods

Field sampling can be performed by several methods. The most commonly accepted methods are line transect sampling and quadrat methods.

• The Downed Woody Material Inventory System, also called Browns method is the most commonly used line transect method for sampling fuels. In this method, several plots are installed at a site. At each plot, three 11 m (35 ft. or 1/2 chain) long, 1.8 m (6 ft) tall planar intercept transects are installed at three azimuths. Along each transect, all 10-hour timelag fuels are counted within 3 m (10 ft) of the endpoint, and 100-hour timelag fuels are counted along the entire 10 m transect. The diameter at the point of intersection, species, and decay status of 1000-hour timelag fuels are measured along the entire 11 m transect as well (Brown 1974). Additional information on this technique is available at the FIREMON website.

• The WWU tree-based transect method (Taylor and Fonda 1990) is a similar method that is applicable to areas with fuels clustered around trees, as in long fire return interval ecosystems (e.g., sand pine scrub) or long-unburned fire-prone forests (fire excluded longleaf pine – hardwood ecosystems).

• Living and dead understory, shrub, and tree fuels can be measured along line transects (as above) or by quadrat methods. In quadrat sampling, herbaceous and forest floor material are harvested from known areas (usually 1 m² for vegetation, smaller for forest floor material), ovendried, and weighed. Shrub and canopy fuels are estimated with using line transects, allometric equations, or using crown densiometers to estimate foliage fuel biomass.

Photo series

A sequence of photos called a "photo series" can provide a quicker and easier means of quantifying fuel loads than field sampling methods, particularly when exact fuel amounts are not required (Reeves 1988, Ottmar and Vihnanek 2000). Photo series consist of a site’s photograph and the fuel loading data associated with the conditions in that specific photograph. Fire managers can then utilize the photographs (often interpolating between more than one) to visually estimate their site’s fuel loading and fire behavior.

Photo series are available for the following southeastern ecosystems:

• Longleaf pine sandhills (Ottmar and Vihnanek 2000, Ottmar et al. 2003)
• loblolly pine, eastern white pine, pitch pine, and Virginia pine (Lynch and Horton 1983)
• poscosins  (Ottmar and Vihnanek 2000)
• freshwater marsh ecosystems (Ottmar and Vihnanek 2000)
• southern pine: post-hurricane (Wade et al. 1993)
• hardwoods with white pine (Ottmar et al. 2003)
• Southern pine and hardwoods: pre and post fire (Scholl and Waldrop 1999)
• pine-hardwood logging slash (Sanders and Van Lear 1987, 1988)

More information on photo series can also be found at the USDA Forest Service Fire and Enviromental Research Applications team website.

Tabular methods

The following tables can also be used to estimate fuel loadings (tons/acre) on an area.

Litter weights:

• (Total litter weight under shortleaf pine stands as affected by stand basal area and age of rough)

• (Total litter weight under slash and longleaf pine stands as affected by stand basal area and age of rough)

• (Total litter weight under loblolly pine stands as affected by stand basal area and age of rough)

• (Litter and duff weight accumulations in developing undisturbed loblolly pine plantations from 6 to 23 years)

• (Litter buildup in upland hardwood stands )

• (Estimated available litter fuel weights for different total litter weights and moisture contents)

Vegetation:

• (Understory vegetative dry weight in the palmetto-gallberry type as related to age of rough and understory height)

• (Total available grass accumulations under pine overstories)

• (Estimated total available fuel (litter + vegetation) as a function of total litter layer moisture content, total litter weight, and understory vegetative dry weight)

• (Dry weight of foliage and branches to 1-1/2 inches in diameter for individual, shortleaf pine trees )

Slash:

• (Available fuel 1 inch in diameter and less in the unpiled pine logging debris type)

Remote sensing

Recently, remotely sensing has been used for fuel characterization and monitoring. This technique is based on the fact that a high correlation exists between spectral variation in remote sensing imagery and fuel variation. This technique involves a significant amount of data exploration to establish relationships between the imagery and fuel features on the ground. Computer algorithms are then used consistently classify the imagery based on the identified relationships.

Chemical properties of fuels affect their heat content and the types of emissions released during a fire (Hough 1969, Shafizadeh et al. 1977, Rundel 1981, DeBano et al. 1998). Forest fuels are composed of living plants and dead plant parts, both of which are constructed of:

• cellulose (from 35 to 55%): an important source of combustible volatiles and is made up of polymers of hexose sugar D-glucose in a linear structure. Cellulose pyrolizes rapidly, producing much of the gaseous fuel in flaming combustion.

• hemicellulose (15 to 25%): a carbohydrate polysaccharide that is similar to cellulose and is found in the cell walls of many plants.

• phenolics such as lignin and tannins (15 to 30%): lignin is an aromatic polymer that does not degrade as easily as cellulose and hemicellulose. This material gives wood its hardness and structural integrity. If heated to temperatures above 400-450° C, 50% of lignin will volatize while the rest becomes char residue or pyrolizes slowly during smoldering combustion.

• extractives (from 2 to >45% in foliage fuels): extractives consist of organic compounds such as sugars, resins, alcohols, hydrocarbons, and fatty acids. Extractives can be highly flammable and are often the first hydrocarbons to be volatilized (during pre-ignition) and consumed in combustion. Extractives that are volatile compounds, such as oils and resins can increase heat release because of their high energy content.

• mineral ash: minerals (Si, Ca, Mg, K, etc) are found in low concentrations and are not combustible.

The relative proportions of these compounds can affect a fuel’s ignition, combustion, and extinction characteristics through particle-level flammability. Woody fuels are high in cellulose, lignin, and hemicellulose, while foliage contains high amounts of extractives.

Heat content is an important aspect of fuel chemistry influencing fire behavior (Miller 1994). A fuel’s heat content (also termed heat yield, energy content, caloric value, or heat of combustion) is the potential heat energy of the particle or the reaction heat resulting from complete combustion, measured in kilojoules per gram (or KJ/g or Btu/lb). A forest fuel with a higher heat content will burn at a higher temperature and more rapidly than a low heat content fuel.

While standard values of heat contents are often used (18,620 KJ/kg), forest fuels vary in their heat contents (see Table: Plant flammability). The presence of volatile compounds in some live fuels increases heat content, and thus flammability (Miller 1994). For example, resinous pine heartwood has almost twice the heat content of oak wood. Foliage heat content is strongly affected by extractive content, so can vary by species and by season. Heat contents change with age, with some species increasing contents with aging and others decreasing with time (Burgan and Susott 1991, Hough 1969). Heat contents are important inputs into Rothermel’s fire spread equations and several fire models (e.g., BEHAVE).

Adjustments and reductions for heat of combustion

Total, gross, or high heats of combustion describe the caloric content of a fuel as measured by an oxygen bomb calorimeter and expressed as calories per gram of dry fuel weight. The high heat of combustion sets a theoretical maximum on the amount of potential energy available for combustion. The average value for wildand fuels is 4500 cal/g (Pyne 1984). Since ideal burning conditions are seldom approached in the field, the high heat of combustion is usually adjusted or reduced to account for fuel moisture, radiation, and incomplete combustion (Alexander 1982). The reduced value is usually called heat yield. Although heat of combustion is often used interchangeably with heat yield, the heat yield for a particular fuel will vary with the heat of combustion (Pyne 1996).

The first reduction of high heat of combustion is for the latent heat expended in evaporating adsorbed water (Byram 1959). Since this latent heat cannot be spent in pyrolysis and combustion, it reduces the amount of energy returned for the heat invested (Pyne 1984). The high heat of combustion reduced by this standard amount, 1263 kJ/kg, then becomes the low heat of combustion (Alexander 1982). In practice, low heat of combustion varies so little from fuel to fuel (roughly 10%) that a basic value of 18,620 kJ/kg has been used as a constant (Van Wagner 1973, Albini 1976). A second reduction, for fuel moisture content, is 24 kJ/kg per moisture content percentage point (Van Wagner 1972b).

Although Byram (1959) also adjusted heat of combustion for radiation losses in his equation for fireline intensity, there are two arguments against this reduction: 1) there is no sound basis available for estimating radiation heat as a proportion of the total energy output of individual fires of different intensities and, 2) radiation is not really a loss, but contributes greatly to fire behavior (Van Wagner 1973). This reduction is suggested if some special purpose requires an estimate of only convective heat output (Van Wagner 1972b).

Another possible reduction is for incomplete combustion or char formation. Heats of combustion that have been adjusted to account for these heat losses are also called effective heat yields. Effective heat yields can range from 34-78% of high heat yields (Pyne 1984). Since incomplete combustion is so variable and difficult to measure, use of effective heat yields remain a matter of subjective judgement (Alexander 1982).

Fuel moisture content is among the most important fuel characteristics affecting fire behavior (Byram 1959, Pyne et al. 1996). It determines how much fuel is available for burning, and ultimately, how much is consumed (Miller 1994). Moisture absorbs heat released during combustion, making less heat available to preheat fuel particles to ignition (Burgan and Rothermel 1984). By raising a fuel’s heat capacity, fuel moisture content influences ignition. At high moisture contents, the heat required to evaporate moisture in fuels is more than the amount of heat available in the firebrand (Simard 1968 in Miller 1994), and combustion can be stopped. This point is termed the moisture of extinction (also called extinction moisture content). Moisture of extinction is a function of the fuel type. For most dead fuels (forest floor duff is an exception), the moisture of extinction is between 12 and 40 percent. For live fuels the moisture of extinction generally exceeds 120 percent (fuel moisture is expressed per unit of dry fuel weight, making moisture contents >100 percent possible).

Both live and dead fuels can slow, stop, or contribute to fire spread, depending on their moisture content. The factors that regulate fuel moisture differ among live and dead fuels. The primary determinants of live fuel moisture content are: internal factors that regulate diurnal and seasonal changes, climate, site factors that affect the fuel environment, phylogenetic differences among species groups (evergreen vs. deciduous), and differences among plant tissues (leaves vs. stems). Fuel moisture in dead fuels ranges widely based on particle size, short and long-term weather changes, topography, decay class, and fuel composition (Byram 1959). How these factors regulate live fuel moisture is briefly explained in the following sections:

• Factors Regulating Moisture Content of Live Fuels

• Factors Regulating Moisture Content of Dead Fuels

Fuel moisture content is a primary variable in all fire behavior prediction models, affecting ignition, combustion, amount of available fuel, fire severity, and smoke generation. Estimating fuel moisture is therefore a critical step in planning a prescribed fire or assessing risk of fire danger. There are several ways to measure fuel moisture, including volumetric analysis, gravimetry, moisture indicator sticks, formulas and graphs, and ocular methods, explained briefly in the following section:

• Determining Fuel Moisture

Land managers can also obtain quick estimates of live and dead fuel moisture from the Wildland Fire Assessment System which produces daily maps of both live fuel moisture and dead fuel moisture across the U.S.:

• Map of estimated live fuel moisture

• Map of estimated 10- hour fuels

• Map of estimated 100-hour fuels

• Map of estimated 1000-hour fuels

Living plants and dead fuels respond quite differently to weather changes. The moisture content of a living plant is closely related to its physiology. The major variations in moisture are seasonal in nature, although shorter term variations are also brought about by extreme heat and drought. The Wildland Fire Assessment System produces daily maps of live fuel moisture across the U.S.

Internal factors

Fuel moisture fluctuates diurnally, with fluctuations greatest in fine living fuels and least in large living tissues. Fuel moisture also fluctuates seasonally, with young foliage having water contents two to three times the content in senescent foliage.

Climatic variation

Climate not only influences the length of the growing season, but can also cause periods of cold-induced dormancy or heat or drought induced quiescence. Long-term climate patterns (particularly extended droughts) can override weather events, creating a lag in changes in living plants’ fuel moisture contents.

Site factors and fuel environment

Site conditions such as soil and canopy cover can cause differences in moisture content within the same species. Elevation and aspect affect local microclimate and produce local differences in seasonal development of many plant species (Schroeder and Buck 1970). In mountain topography, for example, lower elevations and southern exposures favor the earliest start of the growing season.

Species groups

Seasonal moisture patterns vary among deciduous and evergreen species and woody and herbaceous species. Deciduous species generally have higher moisture content than evergreen species. Moisture content in deciduous species also varies more throughout the season, reaching a peak soon after bud break and decreasing after seasonal growth has finished. Because they retain old leaves for several years, evergreen shrubs have a more complex pattern of seasonal moisture content. As with evergreen shrubs, conifers have complex annual patterns of moisture content because they retain needles for several years. In general, old needles reach their lowest moisture content when new needles are being formed. In the southeastern U.S., conifers may flush more than once during the growing season. Moisture content for herbaceous species may be more variable throughout the year than any other species group. With annual species, all of the plant is new tissue at the beginning of the growing season, and all of the plant can become cured at the end. Once cured, herbaceous species respond to atmospheric conditions as a dead fuel. This is particularly true for grasses in areas with hot, dry summers (Byram 1959, Blackmarr and Flanner 1975).

Plant part

The major live fuel categories are foliage, twigs, branches, stems, cones and fruits, and roots. Foliage and other non-woody fuel moisture content are generally very high (>100 – 300 percent), varying among and within species. Foliage moisture is greatest in new foliage, decreasing as the foliage ages. Southeastern pocosin shrub species foliage emerges with 200 to 300% moisture contents and declines rapidly to 100-150% at the end of the growing season. Retained foliage on evergreen shrubs has 100% moisture content, varying little during the year (Blackmarr and Flanner 1975). Cones and other fruiting structure fuels have moisture contents that vary tremendously by season and developmental stage. Small woody fuels (twigs and branches) have lower fuel moisture contents, but these values are considered stable. Stems of pocosin shrub species maintain moisture contents at 100% throughout the year (Blackmarr and Flanner 1975). Still larger living woody fuels (small stems and boles) have lower, albeit more stable moisture contents.

Evergreens growing in climates having marked seasonal changes generally have seasonal growth cycles. Leaves that have lived through a dormant period increase in moisture content at the beginning of the new season from a minimum of perhaps 80-100 percent to a maximum of perhaps 120 percent within a few weeks. These values are typical, but do not necessarily apply to all species and regions. Moisture decreases slowly after this modest increase until the minimum is again reached at the onset of dormancy.Within a few days of the initial increase in moisture in old leaves, twig and leaf buds open and a new crop of leaves begins to emerge. Their initial moisture may exceed 250 percent. Leaves may emerge quickly, or over an extended period, depending on species and the character of the weather-related growing season. The average moisture content of the new growth drops rapidly to perhaps 150 percent, as the new leaves grow in size until about midsummer, and then more slowly, matching the moisture content of the older foliage near the end of the growing season.

Different species of evergreen trees and shrubs characteristically retain a season's crop of foliage for different periods of years. This may vary among species from one season to five or more. There are also differences within species, due partly to age, health, and stand density, but mostly to the weather-dictated character of the growing season. Thus, in years of poor growth there is normally little leaf fall, and in years of lush growth the fall is heavy. As crown canopies become closed, leaf fall tends to approximate foliage production. The oldest foliage, that closest to the ground, is the first to fall, and, in time, the lower twigs and branches that supported it must also succumb and add to the dead fuel supply.

There are exceptions, of course, to the normal, seasonal growth and leaf-moisture cycle, and to the annual replenishment of foliage. Particularly striking are the variations found in the drought-resistant brush and chaparral species in the semiarid West. It is not uncommon for midseason soil-moisture deficiency to cause cessation of growth in these species, with foliage moisture lowering to between 40 and 50 percent. Usually, these plants retain the ability to recover after the next rain. Prolonged severe drought, however, can prove fatal to major branches or even to whole shrubs. Conflagration potential is then at its peak.

The live foliage of evergreens as a class is usually more combustible than that of deciduous species. There are several reasons, but differences in their moisture regimes are most important. All deciduous foliage is the current year's growth, and it maintains relatively high moisture content during most of the growing season. Evergreens, on the other hand, and particularly those that retain their foliage for a number of years, have much lower average foliage moisture during the growing season. Old-growth foliage with its lower moisture may constitute 80 percent or more of the total evergreen foliage volume. Among the evergreens, too, there is greater tendency toward a mixture with dead foliage, branches, and twigs.

Annual range grasses are much more sensitive to seasonal and short-term weather variations than are most other fuels. These grasses are shallow-rooted and thus depend primarily on adequate surface soil moisture for full top development. At best, annuals have a limited growth season. They mature, produce seed, and begin to cure or dry. But deficient surface moisture at the beginning of the season, or its depletion by hot, dry weather may shorten the growth period. Similarly, because of the weather, the curing time may vary from 3 weeks to 2 months after noticeable yellowing.

Perennial grasses have deeper, stronger root systems than annuals and are somewhat less sensitive to short-term surface soil moisture and temperature changes. In regions that have marked growing seasons limited by hot, dry seasons or cold winters, the perennial grasses have, however, a growth and curing cycle similar to annuals, but dieback affects only leaves and stems down to the root crowns. The principal differences in moisture content result from a later maturing date and a slower rate and longer period of curing. In warm, humid areas, some stems and blades cure and die while others may remain alive, although more or less dormant. Often, such mixtures will burn in dry weather.

Dead fuels absorb moisture through physical contact with water (such as rain and dew) and adsorb water vapor from the atmosphere. The drying of dead fuels is accomplished by evaporation. These drying and wetting processes of dead fuels are such that the moisture content of these fuels is strongly affected by fuel sizes, weather, topography, decay classes, fuel composition, surface coatings, fuel compactness and arrangement:

Particle size

Small fuels can gain and lose moisture faster than large fuels. Based on this principle, dead woody fuels are divided into 4 diameter classes, also called timelag classes: < 1/4, 1/4-1, 1–3, and > 3 in. See Timelag classes.

The Wildland Fire Assessment System produces daily maps of dead fuel moisture across the U.S. based on time-lag classes: Map of estimated 10- hour fuels, Map of estimated 100-hour fuels, and Map of estimated 1000-hour fuels.

Weather

Short-term and seasonal/annual patterns in weather determine fuel moisture changes, and thus fuel consumption. Primary weather factors that determine fuel moisture changes are sun, wind, precipitation, relative humidity, and air temperature (Schroeder and Buck 1970). Sunny skies, elevated wind speeds, lack of precipitation, low humidity, and warm temperatures all act to dry fuels. Cloudy or hazy skies, still winds, precipitation, elevated humidity, and cool temperatures either act to increase fuel moistures or prolong their present moisture state. Long-term events like droughts affect 100- and 1000-hour fuels significantly. For more information see: Effects of weather and topography on fuel moisture.

Topography

Topography affects fuel moisture indirectly by influencing microclimate. Fuel moisture content is generally higher on north facing slopes than south slopes because there is less direct exposure to sun. Fuels are also moister at high elevations due to lower temperatures and higher relative humidity (Pyne et al. 1996). For more information see: Effects of weather and topography on fuel moisture.

Decay class

Moisture can be gained and lost more rapidly in decayed wood because particle size is reduced, surface area-to-volume ratio is increased, and moisture-holding capacity is lost. Dry decayed wood combusts and is consumed rapidly. Decay may also remove flammable volatile compounds from dead fuels.

Fuel composition

The major dead fuel categories are dead standing herbaceous material, leaf litter, cones, fallen twigs and branches, fallen logs, and standing snags. Fuel moisture response varies among dead woody fuels, deciduous leaf litter, grass litter, and coniferous needle litter (see Timelag classes). The moisture content of the forest floor complex also has unique characteristics based on its thickness, composition, arrangement, and compactness. Slash (logging debris) presents a special fuel complex.

Surface covering

A covering of organic material, such as bark, waxes, or cutin, can slow the movement of moisture in or out of dead woody fuels. For example, dead woody fuel with bark can gain and lose moisture at 2/3 the rate of fuels without bark (Simard et al. 1984).

Dead vegetation retains its original structure of cells, intercellular spaces, and capillaries. It can soak up liquid water like a blotter, only more slowly, until all these spaces are filled. Dead vegetation may hold two or more times its own dry weight in water. Fine materials may absorb that much in a matter of minutes, while large logs may require a season or more of heavy precipitation. In some climatic regimes, the centers of large materials may never become completely saturated. One reason is that the rate of penetration slows down with increasing distance from the surface.

A second and equally important consideration in understanding fuel-wetting processes is that the materials making up the dead cell walls are hygroscopic. Hygroscopic materials have an affinity for moisture that makes it possible for them to absorb water vapor from the air. This process is one of chemical bonding.

Molecules of water are attracted to, penetrate, and are held to the cell, fiber, or walls by the hygroscopic character of the cell material. The water molecules that penetrate and the few molecular layers that adhere to the cell walls are called bound water. The hygroscopic bond between the cell walls and the water molecules is strong enough to effectively reduce the vapor pressure of the bound water. The layer of water molecules immediately in contact with a cell wall has the strongest hygroscopic bond and lowest vapor pressure. Successive molecular layers have progressively weaker bonds until the cell walls become saturated. At that point, the vapor pressure in the outer layer of water on the cell wall is equal to that of free water, or saturation pressure. The amount of bound water at the fiber-saturation point varies with different materials. For most plant fuel it is in the range of 30 to 35 percent of the fuel dry weight.

The result of the bonding phenomenon is that free water cannot persist in a cell until the cell walls become saturated. Then free water can pass through the cell walls by osmosis. Below the saturation level, moisture is evaporated from cell walls of higher moisture content and taken up by cell walls of lower moisture content until the moisture in each cell attains the same vapor pressure. In this manner, much of the moisture transfer within fuels is in the vapor phase and always in the direction of equalizing the moisture throughout a particular piece of fuel.

Dead fuels will extract water vapor from the atmosphere whenever the vapor pressure of the outer surface of the bound water is lower than the surrounding vapor pressure. In a saturated atmosphere, this may continue up to the fiber-saturation point. Full fiber saturation rarely persists long enough in the absence of liquid water to permit the necessary internal vapor transfer.

See also: Fuel-drying Processes.

Fuel drying is accomplished only by evaporation to the atmosphere. The moisture content of dead fuels thoroughly wetted with free water within and on the surface decreases in three steps in a drying atmosphere, with different drying mechanisms dominant in each period:

1. Constant-rate period: The rate here is independent of both the actual moisture content and the hygroscopic nature of the fuel. It ends at the critical moisture content, the condition in which the total fuel surface is no longer at or above fiber saturation.

2. Decreasing-rate period: During this period, there is a decreasing saturated fuel surface area and an increasing proportion of moisture loss through the slower removal of bound water. The period ends when all the fuel surface reaches the fiber-saturation level.

3. Falling-rate period: During this period, the hygroscopic nature of dead fuel becomes dominant in the drying process.

The process of moisture loss in the constant-rate period is somewhat simpler than those of the succeeding steps. Drying takes place by evaporation exactly as from any freewater surface. It will proceed whenever the surrounding vapor pressure is less than saturation pressure, and at a rate generally proportional to the outward vapor-pressure gradient. Wind speed during this period does not affect ultimate attainment of the critical moisture content level. But it does affect the time required to reach that point. When there is evaporation from a water surface in calm air, a thin layer next to the interface between the free water and air tends to become saturated with water vapor. This saturation near the water surface decreases the evaporation rate and dissipates only by relatively slow molecular diffusion in the air. Wind breaks up this thin layer and blows it away, thereby speeding up the evaporation process.

The intermediate decreasing-rate period may best be described as a transition step in which there is a variable change in moisture loss rate. This rate begins changing slowly within the defined limits from the linear rate of the constant-rate period to the orderly decreasing rate characteristic of the falling-rate period. Variations in the rate of drying during the decreasing-rate period are caused by fuel and environmental factors that are difficult to evaluate and for which no general rules are available. This period is often considered as part of what we have called the falling-rate period when the error involved in calculations is considered tolerable. It is separated for our purposes because it applies only to drying and is not reversible in the sense of vapor exchange between fuel and air as is the case in the falling-rate period. Wind speed still plays a significant role in the drying process during this period.

The falling-rate period of drying depends upon an outward gradient between the bound-water vapor pressure and the ambient vapor pressure in the atmosphere. As moisture removal progresses below the fiber-saturation point, the bound-water vapor pressure gradually declines, and the vapor-pressure gradient is gradually reduced. Either of two conditions must prevail to assure continued significant drying: One is to maintain a surrounding vapor pressure appreciably below the declining bound-water vapor pressure; the other is addition of heat to the fuel at a rate that will increase its temperature and correspondingly its bound-water vapor pressure. Both processes operate in nature, sometimes augmenting and sometimes opposing each other (see Moisture Equilibrium).

As drying progresses toward lower moisture-content values, a vapor pressure gradient is established within the fuel. The external vapor pressure needed to maintain this gradient must therefore be quite low. Under these conditions, molecular diffusion into the atmosphere is more rapid than that within the fuel. This results in a lesser and lesser tendency for thin layers of higher vapor pressure to form at the fuel surface. For this reason, the effect of wind speed on drying gradually decreases at moisture levels progressively below fiber saturation. The effect may never be eliminated, but at low moisture levels it has little practical significance.

See also: Fuel-wetting Processes.

Dead fuels are categorized into fuel diameter classes named according to the timelag principle (Pyne et al. 1996). This principal is based on the fact that the proportion of a fuel particle exposed to weather is related to its size. Small diameter fuels can change rapidly in response to weather changes, while larger diameter fuels are slower to respond. A timelag is the time required for a fuel particle to reach 63% of the difference between the initial moisture content and the equilibrium moisture content (or equilibrium with changed atmospheric conditions). The categories are named for the “midpoint” of the response time of each fuel category: 1-hour fuels respond in less than 2 hours, 10-hour fuels respond in 2 to 20 hours, 100-hour fuels respond in 20 to 200 hours, and 1,000 hour fuels respond in greater than 200 hours. Below are typical fuels and fire behavior for each of these 4 time lag classes.

The Wildland Fire Assessment System produces daily maps of dead fuel moisture across the U.S. based on time-lag classes: Map of estimated 10- hour fuels, Map of estimated 100-hour fuels, and Map of estimated 1000-hour fuels.

1-hour time lag fuels (< 0.625 cm (0.25 in.) diameter)

1-hour time lag fuels are the most important for carrying surface fires and their moisture content governs fire behavior. One-hour fuels include fallen needle and leaf litter, grassy fuels, lichens, and small twigs. Within this category, response times vary by fuel type. Lichen, grass, and well-cured needles respond to changes faster than freshly fallen needles and hardwood leaves. Due to their high surface area to volume, low moisture content, and location in the combustion zone, they produce little smoke and have low flame residence time. One-hour fuels are consumed by both flaming and smoldering combustion, regularly undergoing complete consumption in most surface fires.

10-hour timelag fuels (0.625 - 2.5 cm (0.25 to 1 in.) diameter)

Common 10-hour fuels include small branches and woody stems. Due to their resistance to drying and greater heat capacity, 10-hour fuels often do not combust in low-intensity surface fires. When moisture is low, however, 10-hour fuels can carry hot fires and help ignite larger (100- and 1000-hour) fuels. Ten-hour fuels are readily consumed when fuel moistures are low.

100-hour timelag fuels (2.5 cm - 7.6 cm (1 - 3 in.) diameter)

Larger downed woody debris is common 100-hour forest fuels. These fuels take longer to dry, deterring their consumption under most conditions. Likewise, 100-hour fuels are slow to gain moisture, so they can combust after prolonged drought, even with recent precipitation. When 100-hour fuels ignite they can burn for hours, in mixtures of flaming and smoldering combustion. Decay of 100-hour fuels can alter their response and makes them combust more readily than intact fuels.

1000-hour timelag fuels (> 7.6 cm (3 in.) diameter)

These fuels, which include large downed branches, logs, and tree stumps, burn only under prolonged dry conditions, or when sufficiently pre-heated by adjacent fuels. Since they do not commonly burn, 1000-hour fuels can act as firebreaks and cause fire shadows. When they do burn, 1000-hour fuels are common smoldering fuels and can burn for days after ignition, creating air quality and re-burn hazards.

One method of expressing absorption and drying rates based on both equilibrium moisture content and fuel characteristics makes use of the timelag principle. According to this principle, the approach to equilibrium values from moisture contents either above or below equilibrium follows a logarithmic rather than a straight-line path as long as liquid water is not present on the surface of the fuels.

If a fuel is exposed in an atmosphere of constant temperature and humidity, the time required for it to reach equilibrium may be divided into periods in which the moisture change will be the fraction (1-1/e) ≈ 0.63 of the departure from equilibrium. The symbol, e, is the base of natural logarithms, 2.7183. Under standard conditions, defined as constant 80° F temperature and 20 percent relative humidity, the duration of these time periods is a property of the fuel and is referred to as the timelag period. Although the successive timelag periods for a particular fuel are not exactly equal, the timelag principle is a useful method of expressing fuel-moisture responses if average timelag periods are used.

To illustrate the moisture response, let us assume that a fuel with a moisture content of 28 percent is exposed in an environment in which the equilibrium moisture content is 5.5 percent. The difference is 22.5 percent. At the end of the first timelag period, this difference would be reduced 0.63 x 22.5, or about 14.2 percent. The moisture content of this fuel would then be 28 -14.2, or 13.8 percent. Similarly, at the end of the second timelag period the moisture content would be reduced to about 8.6 percent, and so on. The moisture content at the end of five or six timelag periods very closely approximates the equilibrium moisture content.

The average timelag period varies with the size and other factors of fuels. For extremely fine fuels the average period may be a matter of minutes, while for logs it ranges upward to many days. Using the timelag principle, we can describe various fuels--irrespective of type, weight, size, shape, compactness, or other physical feature--as having an average timelag period of 1 hour, 2 days, 30 days, and so on. Dead branchwood 2 inches in diameter, for example, has an average timelag period of about 4 days. Logs 6 inches in diameter have an average timelag period of about 36 days. A 2-inch litter bed with an average timelag period of 2 days can be considered the equivalent, in moisture response characteristics, of dead branchwood (about 1.4 inches in diameter) having a similar timelag period if there is no significant moisture exchange between the litter and the soil (see Timelag classes).

Moisture equilibrium has meaningful application to forest-fuel moisture only in the range of moisture-content values between about 2 percent and fiber saturation. This is the range covered by the falling-rate period of drying. Fuel will either gain or lose moisture within this range according to the relative states of the fuel and its environment. The amount, rate, and direction of moisture exchange depend on the gradient between the vapor pressure of the bound water and the vapor pressure in the surrounding air. If there is no gradient, there is no net exchange, and a state of equilibrium exists.

The equilibrium moisture content may be defined as the value that the actual moisture content approaches if the fuel is exposed to constant atmospheric conditions of temperature and humidity for an infinite length of time. The atmospheric vapor pressure is dependent upon the temperature and moisture content of the air. The vapor pressure of the bound water in fuel depends upon the fuel temperature and moisture content.

Assuming that the fuel and the atmosphere are at the same temperature, then for any combination of temperature and humidity there is an equilibrium fuel-moisture content. At this value, the atmospheric vapor pressure and the vapor pressure of the bound water are in equilibrium. This point almost, but not quite, exists in nature. Small vapor-pressure differences can and do exist without further moisture exchange. This is demonstrated by the fact that a dry fuel in a more moist environment reaches equilibrium at a lower value than a moist fuel approaching the same equilibrium point from above. For this reason also, reduction of humidity to zero does not reduce fuel moisture to that value. Vapor exchange involving bound water is not as readily attained as is free water and atmospheric vapor exchange. At low vapor-pressure gradients involving bound water, there is not sufficient energy at normal temperatures and pressures to eliminate these small gradients.

Equilibrium moisture content has been determined in the laboratory for numerous hygroscopic materials, including a variety of forest fuels. The usual procedure is to place the material in an environment of constant temperature and humidity, leaving it there until the moisture content approaches a constant value. The process is then repeated over the common ranges of humidity and temperature encountered in nature. Continuous or periodic weighing shows the changing rates at which equilibrium is approached from both directions. Different fuel types usually have different equilibrium moisture contents, but for most fire-weather purposes it is satisfactory to use the average determined for a number of fuels.

The rates at which moisture content approaches the equilibrium value vary not only with the kind of fuel material, but with other characteristics such as fuel size and shape, and the compactness or degree of aeration of a mass of fuel particles. For any one fuel particle with a moisture content below fiber saturation, the rate of wetting or drying by vapor exchange is theoretically proportional to the difference between the actual moisture content and the equilibrium moisture content for the current environmental conditions.

This means, for example, that when actual fuel moisture is 10 percent from its equilibrium value, the rate of increase or decrease is 10 times as rapid as if the moisture were within 1 percent equilibrium. This relationship indicates that moisture content approaching equilibrium follows an inverse logarithmic path.

Use of the equilibrium moisture-content concept makes it possible to estimate whether fuel moisture is increasing or decreasing under a particular environmental situation, and the relative moisture stress in the direction of equilibrium. This by itself, however, is a poor indicator of the quantitative rate of moisture-content change. To it, we must also add the effect of size or thickness of the fuel in question. Applying the time-lag principle allows us to divide fuels into time-lag classes based on their size and thickness.

The moisture content of live and dead vegetation is a product of the cumulative effects of past and present weather events. Fuel moisture changes as weather conditions change, both seasonally and during shorter time periods. This fact, coupled with known attributes of different fuels, provides a useful basis for estimating fire potential in any forest or range area. Fuel moisture content limits fire propagation. When moisture content is high, fires are difficult to ignite, and burn poorly if at all. With little moisture in the fuel, fires start easily, and wind and other driving forces may cause rapid and intense fire spread. Successful fire-control operations depend upon accurate information on current fuel moisture and reliable prediction of its changes.

Living and dead fuels have different water-retention mechanisms and different responses to weather. Live fuel moisture is closely related to its physiology. The major variations in moisture are seasonalheat and drought.

Dead fuels absorb moisture through physical contact with liquid water such as rain and dew and adsorb water vapor from the atmosphere. The drying of dead fuels is accomplished by evaporation. The nature of the drying and wetting processes of dead fuels is such that dead fuel moisture is strongly affected by weather elements such as precipitation, air moisture, air and surface temperatures, wind, and cloudiness. Dead fuel moisture contents are also influenced by fuel factors such as surface-to-volume ratio, compactness, and arrangement.

During clear weather, fuel-bed surfaces exposed to full midday sun may reach temperatures as high as 160° F. or more. Not only does this greatly increase the bound-water vapor pressure, but it also warms the air near the surface and reduces relative humidity. The combination often results in surface fuel moistures 4 to 8 percent below those in adjacent shaded areas. Similarly at night, cooling of these exposed fuel surfaces may cause dew to form on them, while it does not form under the tree canopy. Surface fuel moistures and accompanying changes in moisture gradients are thus commonly much greater, and at the same time much more spotty, in open forest stands than under forests having closed-crown canopies. Clouds also tend to reduce the diurnal extremes in fuel moisture.

Wind can increase drying processes by moving moist air away from fuel surfaces. But wind can also have the opposite effect. Moderate or strong winds may affect surface temperatures of fuels in the open and thereby influence surface fuel moisture. During daytime heating, wind may replace the warm air layers immediately adjacent to fuel surfaces with cooler air. This in turn raises the relative humidity in that area and lowers the fuel-surface temperature. Fuel drying is thereby reduced. At night, turbulent mixing may prevent surface air temperatures from reaching the dew point, thus restricting the increase of surface fuel moisture.

North-facing slopes do not receive as intense surface heating as level ground and south exposures, so they do not reach the same minimum daytime moistures. The highest temperatures and lowest fuel moistures are usually found on southwest slopes in the afternoon. In mountain topography, night temperatures above the nighttime inversion level ordinarily do not cool to the dew point; therefore, surface fuel moistures do not become as high as those at lower elevations.

The moisture content of ground fuels is influenced by three different moisture gradients, or layers of differing water potential: one between the fuel and the air, another between the fuel and the soil, and still another between the top and bottom of the fuel bed itself. Only the upper surface is exposed to the free air while the lower surface is in contact with the soil. In deep and compact fuel beds, air circulation in the lower layers may be nearly nonexistent.

Precipitation soaking down through the fuel into the soil may produce relative humidities near 100 percent at the lower levels, and this can persist for appreciable times. Subsequent drying starts at the top and works downward. In deep fuels, it is not uncommon for the surface layer to become quite flammable while lower layers are still soaking wet.

Reverse gradients also occur after prolonged drying, resulting in the topsoil and lower duff becoming powder dry. Then morning dew on the surface, high relative humidity, or a light shower may cause a downward moisture gradient.

These changes in upward and downward moisture gradients are common in most compacted fuel beds. In some situations, they may even be part of the diurnal cycle of moisture change in response to diurnal changes in temperature and relative humidity. This is particularly true in open forest stands where much of the surface litter is exposed to direct solar heating during midday and to direct radiant cooling to the sky at night.

The amount of fuel available for combustion is often determined by these interior moisture gradients. In some cases, for example, fire may only skim lightly over the surface; in others, the entire dead-fuel volume may contribute to the total heat output of the fire.

Slash (trunks, branches, and tree tops) from thinning or harvest cutting of coniferous forests is a special and often particularly hazardous kind of dead fuel. Often, it is flammable from the time it is cut, but it is particularly hazardous if added to significant quantities of flammable dead fuels already on the ground. As the slash dries, it becomes more and more flammable. The slash of different species dries at different rates, and within species the drying rates depend on degree of shading, season of cutting, weather, and size of material. Needles and twigs dry faster on lopped than on unlopped slash. Within a matter of weeks, however, it is not necessary to consider slash needle and twig moisture different from that of older dead fuels. Stems, of course, require longer periods of seasoning to approach the fuel moisture of their older counterparts.

Fuel moisture content can be determined in several ways, from cheap and approximate to expensive and precise. Methods discussed here by no means include all field or laboratory methods, but do include: titrimetry (volumetric analysis), gravimetry, moisture indicator sticks, solvent distillation, statistical correlations, formulas and graphs, ocular methods, and remote sensing.

Titrimetry

Titrimetry or volumetric water content measures percent water content of a known volume of fuel. This analysis is performed by:

• Measuring the dimensions of the fuel volume (the length, width, and height of a fuel volume, i.e., 1 m³) of interest. This known volume is collected and weighed as a “wet” weight.

• The wet sample is then sealed and transported to a drying oven and dried (usually between 70 and 100^oC) to a constant weight.

• The “dry” sample is weighed and subtracted from the “wet” field weight.

• This difference is divided by the original wet weight and the resulting value is the volumetric water content (%).

The advantages of this method are its ability to integrate different fuels for a better estimate of a stand or site fuel moisture and its low cost (Pyne et al. 1996).

Gravimetry

Gravimetric fuel moisture content measures the percent water content of a known weight of fuel. Sampling is very similar to that used for volumetric water content, except that the original volume of fuel does not need to be measured:

• Collecting field fuels and weighing their “wet” weight.

• The wet sample is then sealed and transported to a drying oven and dried (usually between 70 and 100^oC) to a constant weight.

• The “dry” sample is weighed and subtracted from the “wet” field weight.

• This difference is divided by the original wet weight and the resulting value is gravimetric fuel moisture content (%).

The advantages of gravimetric analysis are its ease of sampling (1-, 10-, 100-, and 100-hour fuels can be “grabbed” and weighed without measuring a fuel’s dimensions), low cost, and subjectivity of fuels sampled (Pyne et al. 1996). It is often used to measure moisture content for specific classes of fuels: grass, forbs, woody debris, etc...

Moisture indicator sticks

Moisture indicator sticks are another method of determining fuel moisture. The sticks are 0.625 cm (¼”) diameter x 50 cm (20 in.) long wooden (pine sapwood) dowels that mimic water absorption of 10-hour timelag fuels. Moisture sticks are placed on support brackets 25 cm (10 in.) above a fuelbed and allowed to equilibrate for some time prior to observation. Measurements of moisture are made on a standard scale and applied as a site average. Moisture sticks have the advantage of being site specific and quickly measured.

Tabular, Statistical correlations, and graphs

Tabular or statistical correlation and graphical methods are a popular method of fuel moisture determination. Many tables have been generated that integrate relative humidity, air temperature, and wind speed to yield fine fuel (1-hour timelag fuel) moisture content. These measures are easily acquired (through field, belt weather kit, or weather station measurements) and tables are quick. This method has utility for fine fuels and for occasions when quick determination of fuel moisture is critical. For example, see the following tables to predict litter moisture.

• (Fine fuel moisture content of dead grass (1-hour timelag))

• (Moisture content of pine needle litter)

• (Average moisture of hardwood litter under an upland oak stand in Eastern Tennessee)

• (Total litter moisture content from duration of (last) rain and days since (last) rain data)

Ocular methods

Ocular methods of determining fuel moisture are another method of fuel moisture determination. Ocular methods include observation of vegetation color, either directly or by using remote sensing imagery (Pyne et al. 1996). Experienced burners in the Southeast can manually bend, twist, and break litter fuels (1-hour timelags) to estimate relative fuel moisture contents. These methods can be inexpensive (even analysis of remote sensing data is inexpensive over large landscapes), but require extensive experience and can lack utility in fire management situations.

Solvent distillation

Solvent distillation is another method of determining fuel moisture content. Distillation consists of heating or boiling a fuel sample of known weight or volume. The volatilized gases and vapor are captured, cooled, and collected. The weight or volume of the collected liquid is subtracted from the original weight or volume, and then divided by the original weight or volume. The resulting value is the distillation fuel water content. This method is precise, but is labor and equipment intensive.

Instrumental techniques

Several fuel moisture determining instruments are used in southeastern fuels. Some instruments, developed for lumber analysis, use electric resistance probes to determine woody moisture content. Other techniques use small hand-held heating chambers that vaporize water of fine and small woody fuels. Both of these instrumental techniques have cost and applicability constraints, but there is great potential for hand held, low-cost, precise field fuel moisture devices.

Remote sensing

Recently, remotely sensing has been used to characterize fuel moisture. The Wildland Fire Assessment System produces daily maps of live fuel moisture estimates. The moisture image will be most useful in shrub communities where users have measured live fuels (ground truthed) and developed correlations between measured and image moisture values.

Often, total fuel loads do not equal fuel available for consumption. Fuel availability is dependent on:

• fuel moisture content
• fuel location (ground, surface, ladder, canopy),
• fuel physical properties (continuity, compactness), and
• fuel chemical properties (volatile compounds).

While dead fuels are generally the most important source of available fuels, in some areas of the southeastern US live fuels can comprise a very significant portion of available fuels. Fuel consumption models are available for predicting fuel consumption for hazard reduction, smoke management, and ecological restoration.

Consumption of dead fuels

In terms of fuel consumption, dead fuels comprise the dominant portion of available fuel loads. Dead fuels generally have low moisture contents, are close to flaming and smoldering zones, have smaller particle sizes and altered chemical content than they had when living. Consumption of dead fuels is a major goal for forest managers attempting hazard reduction and ecological restoration throughout southeastern forests.

In most southeastern communities, dead standing herbaceous fuels comprise the most important component of the fuel matrix (see fuel models 1-3). The dead component of bunch grasses (notably wiregrass (Aristida stricta) and bluestems (Schizachyrium and Andropogon spp.)) can be the largest source of fine fuels in pine savannas. Dead standing forbs are also an important source of fine dead fuels, but generally comprise less of the fuel load than dead grasses. Fuel moisture in standing dead fuels is based on short and long-term weather changes, topography, and fuel bulk density (see: Moisture Content of Dead Fuels).

Dead woody fuels are divided into 4 broad categories based on fuel particle size and related drying times: 1-, 10-, 100-, and 1000-hour fuels (see Timelag classes). Fuel moisture in dead fuels ranges widely based on these particle sizes, short and long-term weather changes, topography, decay class, and fuel composition (see: Moisture Content of Dead Fuels). Standing dead hardwood sprouts and saplings killed in previous burns can be an important component of dead woody fuels. In the southeastern U.S., broadscale herbicide application results in a large amount of dead standing 1-hour and 10-hour fuels.

The litter layer is a third source of dead fuel available for consumption. Litter (Oi) horizons contain primarily 1-hour litter fuels that dry quickly and burn with high intensity and short duration, with complete consumption except under elevated fuel moistures. Within or residing on litter fuels is a patchy complex of downed woody 10-, 100-, and 1000-hour fuels which can also alter consumption patterns (see Surface Fuels).

Litter forms a part of the forest floor complex, a matrix of litter, duff, and down woody debris above the mineral soil. Fuel availablity of the forest floor is complex. Each forest floor horizons has different moisture relationships, particle sizes, chemical composition, densities, and depths. Forest floor fuel availability is determined primarily by fuel moisture content and fuelbed bulk density. Separation of available and unavailable fuel is made on depth to moisture, with all dry fuel included as surface fuel and the remaining wet included as ground fuel.

Beneath the litter horizon in many unburned stands is a well-developed duff layer comprised of the fermentation (Oe) and humic (Oa) horizons. Due to decomposition processes, the shade provided by the overlying litter, and its greater packing, the fermentation horizon often retains moisture and resists consumption except during smoldering-phase combustion. Similarly, the humic horizon often has greater moisture content than the overlying litter, but can undergo drying from the mineral soil beneath. In areas where duff has accumulated, duff consumption can be of major importance in southern forests (see Ground Fuels). Duff, when sufficiently dried, can smolder for hours to days post-ignition. Duff smoldering kills surficial tree and plant roots, can kill tree vascular tissues, and with optimal conditions can re-ignite unburned areas (Frandsen 1987, Miyanishi 2001).

Dead standing trees, also called snags, also represent a major dead fuel available for consumption. Snags can ignite during both surface and ground fires and flame or smolder long after the ground or surface fires are extinct. Pine snags often contain resinous heartwood cores that maintain structural integrity even when large portions of the stem are consumed by fire. Standing, burning snags are a source of firebrands if unburned fuels are nearby (e.g., along a fire break). When burning snags fall, they can re-ignite forest floor fuels and contribute to fuel consumption, re-burn hazards, and generate smoke that is a source of both safety and health concerns. Low stumps can also ignite when low moisture conditions prevail. Stumps can smolder both above and below ground, igniting dead roots and posing both re-burn and smoke hazards. Hurricanes, beetle infestations, drought, and intense wildfire can all dramatically increase the number of standing snags and down woody debris in southeastern forests.

Consumption of live fuels

Live fuels differ from dead fuels primarily in their high and stable moisture content, elevation above the surface, and protective structures. In many situations, these factors make live fuels less available for consumption than dead fuels except under extreme events (erratic fire behavior, extended droughts). However in the southeast U.S., live fuels make up an important part of the fuel complex and in fact, can comprise the majority of total available fuel (e.g., in pine flatwoods, herbaceous wetlands, Florida scrub).

The primary determinants of live fuel moisture content are: internal factors that regulate diurnal and seasonal changes, climate, site factors that affect the fuel environment, phylogenetic differences among species groups (evergreen vs. deciduous), and differences among plant tissues (leaves vs. stems). How these factors regulate live fuel moisture is briefly explained in Live Fuel Moisture.

Live fuel types can be divided into plant leaves, stems, twigs, branches, cones, flowers, and roots. Most important for consumption are generally leaves, small woody fuels (twigs and branches), and roots. Plant leaves have high surface area-to-volume ratios, lower fuel moisture contents, and are often consumed in surface fires. Many leaves contain volatile compounds that enhance flammability and often carry surface fires. Examples of this phenomenon are prominent in gallberry (Ilex glabra), wax-myrtle (Myrica cerifera) and saw palmetto (Serenoa repens) of pine flatwoods ecosystems. Small woody structures (twigs and branches) can be consumed when located on the forest floor or in shrubs. Small-diameter living plant roots are consumed in ground fires in organic soil and forest floor. Ladder fuels can be consumed and encourage consumption in lower branches and tree canopies. Limited consumption occurs in crown fires; when this does occur, fires are extremely severe, with extreme fire behavior and excessive tree mortality resulting.

For a discussion of the species-specific factors that influence flammability at an individual plant level, see Plant Flammability.

It is of great interest and utility to predict fuel consumption for hazard reduction, smoke management, and ecological restoration. The primary fuel consumption model is CONSUME developed by the USDA Forest Service for many fuels, including southeastern forest fuels. CONSUME utilizes woody loading, forest floor bulk density and loading, shrub and grass cover, and overstory density to model consumption and emissions under various fuel moistures.

The following discussion relates to species-specific factors that control the flammability of live plants. For a discussion of factors that control overall fuel availability of both live and dead plants, see Fuel Availability and Consumption.

Flammability describes how easily a plant will ignite and sustain a flame, how fast it will burn, and how much of it will be consumed in the process. Flammability is a qualitative rather than quantitative measure, i.e., it has no units of measure and therefore is useful in a relative sense for comparing species. Species-specific chemical and structural characteristics vary at the particle, branch, and plant level to affect overall flammability. The characteristics of plants that have the greatest influence on flammability appear to be the quantity of foliage and fine fuels on a plant and the moisture content of these fuels (Etlinger 2000). However, plants with high concentration of organic volatiles can be highly flammable even with high moisture content (Rothermel 1976, Shafizadeh et al. 1977). Flammability of a plant is also influenced by external characteristics such as weather, climate, and location.

There are several plant species in the south that are highly flammable, most notably the southern rough species. Of these, gallberry (Ilex glabra) is the most flammable, due in part to elevated organic volatile content (Shafizadeh et al. 1977, Burgan and Susott 1991). Saw palmetto (Serenoa repens) and wax myrtle (Myrica cerifera) are also highly flammable southern rough species (Shafizadeh et al. 1977, Hough and Albini 1978). Of the southern pines, the long needle pines (longleaf pine (Pinus palustris) and South Florida slash pine (Pinus elliottii var. elliottii) needles are the most flammable (Fonda 2001). Additional information on the flammability of southern plant species is presented in Comparing Flammability of Southern Plants, BROKEN-LINK Table: Flammability characteristics of southern species I, and BROKEN-LINK Table: Flammability characteristics of southern species II.

In addition to providing needed information to Wildland-Urban Interface (WUI) residents engaging in firewise landscaping, quantification of flammability characteristics can contribute to the development of more ecosystem-specific and WUI-specific fire behavior models (Hough and Albini 1978, Rehm et al. 2002). Considerably more empirical data is needed on the flammability of plants, and in particular, those of southern species.

Flammability was initially defined in three components: ignitability, sustainability, and combustibility (Anderson 1970). The ignitability component is the time until ignition once exposed to a heat source. Sustainability is the stability of burning rate, or the ability to sustain fire once ignited. Combustibility is defined as the rate of burn after ignition. The definition of flammability has since been expanded to include consumability, the proportion of mass or volume consumed by fire (Martin et al. 1994).

Anderson (1970) also related the flammability components of individual plants to fire characteristics at an ecosystem level. Ignitability of individual plants drives the chain of ignition in an ecosystem. Sustainability is related to the rate of fire spread and combustibility to fire intensity. The consumability of individual plants is analogous with the amount of fuel available for fire consumption on the ecosystem level (Martin et al. 1994).

Particle level flammability describes the intrinsic components of flammability without the influence of branch or plant structure. Major influences on particle flammability are

• moisture content,
• percent cellulose, hemicellulose, and lignin,
• volatile concentration, and
• silica-free mineral content.

Both extractive and mineral content of plants are species dependent, making these intrinsic flammability characteristics vary significantly among species. How these components influence particle flammability is explained below, followed by a short explanation of how particle flammability is measured.

Moisture content

Moisture content is the most important factor influencing particle flammability. Moisture content was highly significant in the ignition time, maximum burning rate, period of flaming combustion, and flame length of leaf material from Themeda australis, Eucalyptus viminalis, and Xanthorrhoea australis (Gill et al. 1978). However, living fuels in the palmetto-gallberry fuel complex will burn at moisture levels of 100% or more, while dead fuels may not burn at moisture levels of 20-30%; similar observations have been made in southwest chaparral ecosystems and coniferous forests throughout the U.S (Rothermel 1976). Such observations demonstrate that other factors can influence the combustion of vegetation beyond moisture content (1976). See also: Fuel Moisture.

Percent cellulose, hemicellulose, and lignin

Due to their different chemical properties, plant flammability can also be related to the proportion of cellulose, hemicellulose, and lignin in plant tissue (Rundel 1981). Lignin is thermally stable, and it volatilizes slowly with increasing temperatures, losing only 50% of weight at 500^oC (Philpot 1970). In comparison, hemicellulose undergoes combustion at 250^oC with complete volatilization at 500^oC; and cellulose undergoes rapid combustion between 300 and 400^oC (Philpot 1970). The combustion characteristics of these elements are affected by the presence of organic volatiles (discussed below), which can be extracted to test their influence on flammability.

Susott (1982) examined 43 samples from different species and locations and found that foliar material combusted more rigorously than woody material. This was attributed to higher lignin-to-cellulose ratio in woody material, as well as higher extractive concentration in foliage (Susott 1982). In a study of Pseudotsuga menziesii, Pinus ponderosa, Populus tremuloides, Ilex glabra, Arctostaphylos totula, and Serenoa repens, ether and benzene-ethanol extractives contributed up to 60% of the heat release of dried, ground foliar samples (Shafizadeh et al. 1977). Concentration of the extractives in tissues was determined to be a useful but not conclusive prediction of heat release (Shafizadeh et al. 1977). Dried foliar samples from species in a flammable fynbos (South African scrubland) ecosystem were found to have higher crude fat content (oils, fats, waxes, and terpenes) and higher energy content than dried foliar samples from species in a non-flammable forest ecosystem (Van Wilgen et al. 1990). Higher crude fat content, lower foliar moisture content, and higher energy content for fynbos species were thought to contribute, along with structural characteristics of the ecosystems, to the differences in ecosystem flammability (Van Wilgen et al. 1990).

Volatile concentration

Collectively, the presence of extractives (flavonoids, waxes, terpenes, oils, and resins) increases ignitability and combustibility. This occurs because they typically undergo combustion at lower temperatures than cellulose and lignin and are highly flammable at high temperatures (Rundel 1981). Shafizadeh et al. (1977) concluded that total extractive content likely affects flammability when it exceeds 25% of oven dry mass. Owens et al. (1998) found that a 1 mg increase of limonene per dried g of foliage in Juniperus ashei, increased flammability by as much as 30%. However, the same study determined bornyl acetate was negatively related to flammability, decreasing flammability by 2% with a 1 mg per dried g of foliage increase, illustrating that not all extractives increase flammability (Owens et al. 1998). See also: Volatile Compounds.

Mineral content

In an early analysis of the impact of mineral content on flammability, Mutch and Philpot (1970) determined that the silica portion of incombustible mineral ash does not influence plant flammability. Further study of the mineral portion of vegetation revealed that an increase in silica-free ash, as percentage of dry weight, decreased maximum combustion rates and increased residues (Philpot 1970). This indicates that the percent of mineral ash, minus silica, in plant tissue decreased the rate at which the tissue combusted.

Measuring particle level flammability

Particle level flammability is determined by reducing plant material, typically leaves, into a fine, uniform substance. Measurements at this level have been made by thermal evolution analysis (Shafizadeh et al. 1977), oxygen bomb calorimetry (Dickinson and Kirkpatrick 1985, Van Wilgen et al. 1990, Rodr\xc3\xadguez-Añ\xc3\xb3n et al. 1995, Núñex-Regueira et al. 2000, Núñex-Regueira et al. 2001, Williamson and Agee 2002), thermogravimetric analysis (Mutch and Philpot 1970, Philpot 1970, Shafizadeh et al. 1977, Gill et al. 1978, Rogers et al. 1986), thermocouple analysis (Owens et al. 1998), and evolved gas analysis (Susott 1982).

Branch-level flammability

Branch flammability is influenced by both particle flammability and the arrangement of particles into leaf and stem structure. Leaf thickness, surface area-to-volume ratio, and particle density affect flammability at the leaf and stem level.

Time until ignition at 750^o C was directly related to leaf thickness of foliar samples from 32 species (Montgomery and Cheo 1971). In the same study, surface area-to-volume ratio was inversely related to ignition time for the same samples (Montgomery and Cheo 1971).

Heat transfer, in the form of radiation, conduction, and convection, is affected by surface area. In fuels with high surface area-to-volume ratios, heat is transferred faster to the interior causing more rapid combustion (Rundel 1981). In addition, fuels with higher surface area-to-volume ratios can exhibit more rapid water loss, indirectly increasing flammability (Rundel 1981). For examples of species with high surface area-to-volumes, see Surface area-to-volume ratios of Southern Species.

The amount of mass per volume of particles, or particle density, also influences heat transfer, thereby affecting flammability (Rundel 1981). Particle density affects the type of ignition, whether spontaneous (indirect heat source) or pilot-ignited (direct heat source). Lower particle density fine fuels are more likely to spontaneously ignite in the absence of a pilot fire (Brown 1970).

In general, a leaf attached to a plant contributes to fire behavior differently than a similar leaf immediately after being dropped from a plant. In a study assessing the use of the limiting oxygen index method in measuring foliar flammability, mature leaves and freshly fallen leaves from 10 tree and shrub species were tested (Mak 1988). Results showed that the freshly fallen leaves required less oxygen to ignite and sustain burning than the mature leaves of the same plant (Mak 1988). It is unclear how the chemical makeup of the leaves differed.

Methods for testing flammability at the leaf or stem level include muffle furnace tests (Montgomery and Cheo 1971), cone calorimetry (White et al. 1996), and the limiting oxygen index method (Mak 1988).

Plant-level flammability

Horizontal and vertical arrangement of leaves and branches on a plant can affect its overall flammability. Of 10 plant-level characteristics, Etlinger (2000) found that mass and moisture content of foliage were the most important predictors of flammability. Additional information on the flammability of southern plant species is presented in Comparing flammability of southern plants.

Measurements of flammability at this scale have been done with an intermediate scale biomass calorimeter with a line burner ignition source, which is able to measure the heat released from the combustion of an entire plant (Etlinger 2000).

Research limitations of plant flammability research

Measurements of flammability are complex, as they include both internal and external properties, and they can be made using a variety of techniques and equipment. Also, flammability characteristics have been studied to different extents by various methods, are not equally important to plant flammability, nor are they all independent of one another (Shafizadeh et al. 1977, Etlinger 2000, Francis 2000). An accepted methodology is necessary for determining the flammability value of plants that incorporates the entire plant structure, but measures only the most influential characteristics (Frommer and Weise 1995, University of California FPL 2001). In addition to providing needed information to WUI residents engaging in firewise landscaping, quantification of flammability characteristics can contribute to the development of more ecosystem-specific and Wildland Urban Interface (WUI)-specific fire behavior models (Hough and Albini 1978, Rehm et al. 2001). To be able to rank the relative flammability of southern species, more empirical data must be collected.

Climate and weather influence flammability characteristics of plants and related fire behavior, primarily through effects on plant moisture (Agee et al. 2002). Because of this, the location of species in terms of geographical region and location in a landscape can influence plant flammability. Time to ignition for Juniperus pinchotii foliage was dependant on the moisture content of the foliage as well as the average daily mean temperature for the month preceding sampling (Bunting et al. 1983). (See also: Factors Regulating Moisture Content of Live Fuels.)

Season and location had a significant effect on the monoterpenoid content, percent burned, caloric content, and percent moisture content of Juniperus ashei foliage (Owens et al 1998). Low temperature volatiles (up to 300^o C) do not fluctuate seasonally in saw palmetto, gallberry, or wax myrtle (Burgan and Susott 1991). However, there is a seasonal trend in high temperature volatiles (500^o C) in saw palmetto, wax myrtle, and especially gallberry (Burgan and Susott 1991). (See also: Volatiles.)

In addition to weather and climate, past disturbances such as fire may also change flammability with a pattern that is predictable with time since disturbance. Fire can affect moisture content, relative basal area growth rate, and carbohydrate concentration of chestnut oak, scarlet oak, and red maple (Rieske et al. 2002). Fire affected the moisture content and crude fiber content of Kalmia latifolia leaves, but not ash content, when compared to unburned plots (Thackston et al. 1982). Percent cover, stem height, leaf area, and specific leaf area of Kalmia angustifolia, an ericaceous shrub of Newfoundland, were dependent upon forest type and disturbance regime (Mallik 1994).

See also: Particle Flammability, Branch- and Plant-level Flammability

Several studies have compared the overall flammability of different plants, but few have included southern species. The information that does exist is summarized below and presented in BROKEN-LINK Table: Flammability characteristics of southern species I, and BROKEN-LINK Table: Flammability characteristics of southern species II. This section concludes with a summary of the limitations of plant flammability research.

Fuel characteristics of saw palmetto (Serenoa repens) and gallberry (Ilex glabra) are the most extensively studied in the southeast. For the chemical characteristics, gallberry appears to be more flammable than saw palmetto (see BROKEN-LINK Table: Energy content of gallberry and saw palmetto). This is confirmed by Burgan and Susott (1991) in a comparison of flammability characteristics for saw palmetto, gallberry, and wax myrtle (Myrica cerifera). Combusted at temperatures ranging from 300^oC to 500^oC, gallberry burned the most intensely, followed by saw palmetto and then wax myrtle (Burgan and Susott 1991). Gallberry has a very high volatile extractive content (Shafizadeh et al. 1977, Burgan and Susott 1991), which contributes greatly to combustion (Shafizadeh et al. 1977). In contrast, saw palmetto has a low volatile extractive content, so combustion can be attributed to unextractable components such as lignin and cellulose (Shafizadeh et al. 1977).

Plant flammability can help determine the general flammability of community types. For example, Behm (in press) found that understory plants in pine flatwoods (saw palmetto, gallberry, and Lyonia lucida) have greater ignitability, sustainability, and combustibility than understory plants in hardwood hammocks (beautyberry, Callicarpa americana). The results of this study and similar studies have important implications for wildfire hazard to wildland-urban interface (WUI) structures near these different forest types.

Bark, terminal branch, and foliage of Melaleuca have greater energy content (5.621, 4.585, and 5.146 kcal∙g-1) than bark, terminal branch, and foliage of Eucalyptus (4.009, 4.497, and 4.894 kcal∙g-1) (Wang and Huffman 1982). Melaleuca and Eucalyptus are exotics from Australia found in South Florida.

Longleaf pine (Pinus palustris) needles burn faster and more intensely than needles from South Florida slash pine (Pinus elliotii var elliotii) although both are highly flammable (Fonda 2001). Sand pine needles had a significantly longer flame time than longleaf or South Florida slash, but were generally less flammable (Fonda 2001). Longleaf pine litter produces slightly more energy per weight than slash pine litter possibly due to lower mineral content and higher surface area-to-volume ratio for longleaf pine litter (Hough and Albini 1978). No more distinctions between pines can be made at this time, although the observed energy content of the evergreen Juniperus ashei (2.6 kcal∙g-1) was much lower than other evergreen species that have been studied (Owens et al. 1998).

Data on other southern species are much more limited. Based on reported specific leaf area (leaf area per mass), red maple foliage is more flammable than chestnut oak followed by scarlet oak (Rieske et al. 2002). However, the information for these species was not gathered in order to quantify flammability and other information on flammability characteristics is not available. In addition, flammability comparisons between species from different studies are difficult due to differences in environmental conditions and horticultural practices. Although leaf area per plant has been measured for mountain laurel (Thackston et al. 1982) and rosebay rhododendron (Starrett et al. 1993), they are not comparable due to different horticultural production of the nursery plants. The lack of studies comparing flammability characteristics among species precludes any further comments concerning other fuel components of the south.

Weather is the state of the atmosphere and is often described in terms of temperature, humidity, stability, pressure, wind speed and direction, clouds and precipitation. The interaction of these weather elements control many aspects of fire behavior. For example, atmospheric moisture directly effects fuel flammability, and, by its relationship to other weather factors has indirect effects on other aspects of fire behavior. General winds and local winds affects wildfire in many ways. Wind carries away moisture-laden air and hastens the drying of forest fuels. Light winds aid certain firebrands in igniting a fire. It aids fire spread by carrying heat and burning embers to new fuels, and by bending the flames closer to the unburned fuels ahead of the fire. The direction of fire spread is determined mostly by the wind direction. Atmospheric stability is closely related to fire behavior. For example, winds tend to be turbulent and gusty when the atmosphere is unstable, and this type of airflow causes fires to behave erratically. Thunderstorms with strong updrafts and downdrafts develop when the atmosphere is unstable and contains sufficient moisture. Their lightning may set wildfires, and their distinctive winds can have adverse effects on fire behavior.

• Air Temperature: Temperature directly influences many other weather elements, such as moisture, stability, and wind speed and direction. This section discusses temperature and heat; the role of solar radiation in creating diurnal and seasonal variations in air temperatures; how topography, water bodies, and vegetation influence air temperatures; and, vertical variations in temperatures (lapse rates and inversions).

• Atmospheric Moisture: Atmospheric moisture has direct effects on the flammability of forest fuels, and, by its relationship to other weather factors, it has indirect effects on other aspects of fire behavior. This section discusses sources of atmospheric moisture; evaporation, dew points, absolute humidity, and relative humidity; how temperature, wind, and other factors influence relative humidity horizontally and vertically; means of measuring relative humidity.

• Atmospheric Stability: Atmospheric stability is the resistance of the atmosphere to vertical motion. This section discusses methods for determining atmospheric stability; various processes that cause air to rise; how and why stability varies diurnally and seasonally; how cloud-types and other visual clues can be used as indicators of stability; and, what subsidence is and conditions that lead to its development.

• General Winds: General winds are produced by the broadscale pressure gradients but may be modified considerably by friction or other topographic effects. This section discusses laminar flow, eddy formation and the conditions that create mechanical and thermal turbulence; winds created by frontal air-masses; effects of mountains on surface winds; and methods used to measure surface winds.

• Convective Winds: Convective winds originate from small-scale pressure gradients produced by local temperature differences and may be strengthened or weakened general winds. This section discusses land and sea breezes, slope and valley winds and their interaction with general winds; whirlwinds and thunderstorm winds.

• Air Masses and Fronts: Air masses have characteristic weather in their source regions. As they leave their source regions, air masses are modified according to the surface over which they travel. In frontal zones, where two or more air masses meet, considerable weather is concentrated. This section discusses the formation and modification of air masses and their associated weather, warm fronts, cold fronts, and stationary fronts.

• Clouds and Precipitation: The amount of precipitation and its seasonal distribution are important factors in controlling the beginning, ending, and severity of local fire seasons. This section discusses saturation, condensation, sublimation, and precipitation processes as well as types of precipitation and clouds.

• Thunderstorms: The heat energy released in condensation provides the energy for thunderstorms and the violent winds associated with them. Thunderstorms are important in fire control because they start fires by lightning, blow them out of control with the downdraft and outflow, or put them out with rain. This section discusses thunderstorm development, types of thunderstorms, lightning, and tornadoes.

• Fire Climate Regions: This section discusses the general temperature, precipitation, and pressure/circulation patterns in North America and how these climate factors are influenced by broad-scale topographical features and oceans. The general fire climate of the Southeastern United States and 14 other regions are summarized.

Air temperatures also vary vertically, generally decreasing with height. Lapse rates describe this rate of temperature change with height. At night or near coastal areas, inversions can form where air temperature will actually increase with height. Lapse rates and inversions are key components in determining atmospheric stability.

• Temperature, Heat, and the Atmosphere
  • Radiation
• Horizontal Variations in Surface Air Temperature
  • Effects of Forests on Surface Temperature
• Diurnal and Seasonal Variations in Surface Air Temperatures
• Lapse Rates and Inversions
  • Night Inversions
  • Marine Inversions

Temperature can be defined as the degree of the hotness or coldness of a substance and reflects the average molecular activity of the substance. Heat energy represents the total molecular energy of a substance and is therefore dependent on both the number of molecules and the degree of molecular activity. If heat is applied to a substance and there is no change in physical structure (such as ice to water or water to water vapor), the molecular activity increases and the temperature rises. Conversely if heat is lost the temperature will drop.

Heat and temperature differ in that heat can be converted to other forms of energy and can be transferred from one substance to another, while temperature has neither capability. Temperature, however, determines the direction of net heat transfer from one substance to another. Heat always flows from the substance with the higher temperature to a substance of lower temperature, and stops flowing when the temperatures are equal. While the amount of heat received by the cooler substance is equal to that lost by the warmer substance, the temperature changes between the two substances are not necessarily equal.

Since different substances have different molecular structures, the same amount of heat applied to equal masses of different substances will cause one substance to get hotter than the other. In other words the two substances have different heat capacities. In the English system of measurement, one Btu (British thermal unit, a measure of heat capacity) is the amount of heat required to raise the temperature of 1 pound of water by 1 degree Fahrenheit. The specific heat of a substance is the ratio of its heat capacity to that of water (water therefore has a specific heat of 1.0). A substance with a specific heat of 0.5 would increase in temperature by 2 degrees F for each Btu of heat added.

With minor exceptions, solids and liquids expand when their molecular activity is increased by heating and contract when molecular activity decreases due to cooling. The amount of expansion/contraction is dependent upon the material. The expansion/contraction of a liquid is used in thermometers to measure changes in temperature. The response of gases to changes in temperature is more complex as temperature changes may change either the volume or pressure of the gas, or both. If volume is held constant, the pressure will increase as the temperature increases.

The atmosphere receives heat energy from the sun through radiation. This heat energy is converted into thermal energy. Since the atmosphere is not confined, atmospheric processes do not occur at constant volume. Either the pressure is constant or both pressure and volume change. If the pressure remains constant, the volume increases as temperature increases and the volume decreases as the temperature decreases. This change in volume brings about significant changes in the density (mass per unit volume) of the gas. Rising temperatures lead to decreased densities.

When the volume of a gas expands due to increased temperature, the gas must perform work on its environment and therefore expend some of its internal (molecular) energy. Decreasing the internal energy of the gas results in a cooling of the gas. Therefore expansion is a cooling process. Conversely, compression (decreasing the volume of the gas) is a warming process as the environment must do work on the gas to compress it, thereby transferring molecular energy to the gas. An increase in molecular energy of the gas leads to an increase in temperature of the gas.

Radiation is the process by which the earth receives heat energy from the sun. The intensity of this solar radiation at the outer limits of the earths atmosphere is quite constant, but the amount that reaches the earths surface is highly variable, depending largely on the amount of clouds in the atmosphere.

The solar radiation that enters the atmosphere is reflected, scattered by molecules in the atmosphere, or absorbed by the atmosphere and the earths surface. On average, 25% of the incoming solar radiation is reflected by clouds with an additional 3% reflecting off the earths surface. Approximately 17% of the solar radiation will be scattered by molecules in the atmosphere (10 of the 17% ends up being absorbed by the surface while the remaining 7% is lost to space). The remaining incoming solar radiation is absorbed by the atmosphere and earths surface. Water vapor, ozone and carbon dioxide each absorb radiation within certain wavelengths, helping contribute to an average of 22% of incoming solar radiation being absorbed by the atmosphere. The remaining incoming solar radiation (33%) is absorbed by the earths surface, which when combined with the fraction absorbed after scattering leads to an average of 43% of the suns radiation is absorbed by the earths surface.

The radiation absorbed by the earths surface and atmosphere is converted to thermal energy, thus warming the substance. While the solar radiation that reaches the surface warms the surface, the earths average temperature does not change as the earth in turn radiates energy back to the atmosphere and space. The radiation emitted by the earth is very different from that emitted by the sun. Due to the much lower temperature the radiation is of a much longer wavelength which changes how the radiation passes through the atmosphere. Water vapor in the atmosphere is very effective at absorbing this longwave radiation and reradiating a fraction of it back toward the surface, slowing the net energy loss from the surface. This is why cloudy nights are typically warmer than nights with clear skies.

The net balance between incoming solar radiation and outgoing longwave radiation from the earth is responsible for controlling the timing and magnitude of daily high and low temperatures. The earth is continually trying to cool itself by emitting longwave radiation. Beginning at sunrise, solar radiation tries to warm the earth, with this heating peaking at noon and ending at sunset. The timing of maximum temperature does not coincide with the peak in incoming solar radiation as the earth will warm as long as there is a net gain in heat. Likewise the earth will continue to cool as long as there is a net loss of heat, leading to low temperatures occurring near sunrise.

The amount of solar radiation received by any point of the earth is greatly influenced by the tilt of the earths axis and its annual progression around the sun that create the different seasons. During the northern hemisphere summer, the northern hemisphere is tilted toward the sun, allowing the incoming solar radiation to strike the earths surface at a higher (more perpendicular) angle. At this time the southern hemisphere is tilted away from the sun causing the solar radiation to the earth at a lower angle. This lower angle diffuses the radiation over a larger area making it less efficient in heating the surface.

Surface air temperatures are strongly influenced by the differing materials that make up the earths surface as the air is primarily heated/cooled by contact with these surfaces. With the exception of water and ice, most materials that make up the earths surface have a much greater temperature range than does the air. To understand the variability of surface air temperatures, one must understand the factors influencing the temperature of the earths surface and since the majority of heat energy for the earths surface results from solar radiation, one must understand the factors that lead to variations in solar radiation reaching the earths surface.

The key factor influencing the intensity of solar radiation at the earths surface is the incidence angle of the radiation. The closer to perpendicular this angle is the more intense the radiation. Latitude and season have a pronounced large-scale influence on incidence angle, but there are also local factors such as topography to consider. The slope and aspect of the topography can significantly alter the amount of solar radiation reaching the earths surface. As the sun moves across the sky, its rays are more nearly perpendicular to different slopes and aspects at different times of the day. In the northern hemisphere, south-facing slopes receive more direct radiation than do north-facing slopes. Slope and aspect also effect timing of maximum temperature as different aspects reach their peak heating at different times. East-facing slopes reach their max temperature far earlier than those slopes facing west. Surface shading is also a factor consider in assessing local variations in temperature. Shading and scattering of radiation by any means, such as clouds, smoke, or haze in the air and objects such as vegetation, reduce the radiation reaching the ground.

The properties of the surface material also have a significant impact on surface air temperature. Dark materials generally absorb most of the radiation in the visible wavelengths, where as light materials reflect more of these wavelengths. The more radiation absorbed by a material the hotter it will get, therefore dark materials typically heat up more than light colored materials. Transparency is another factor affecting temperature. Water is fairly transparent to solar radiation. This transparency acts to spread the incoming solar radiation throughout a deeper layer of the material, slowing its surface temperature increase. Opaque substances such as soil receive all of the incoming solar radiation at their surface which more effectively heats that surface. A third surface property is conductivity, the ability of a material to transfer heat throughout the material. A material that is a good conductor will rapidly transfer heat energy from the surface to the rest of the material. A poor conductor will end up with a high concentration of heat energy at the materials surface leading to a very high surface temperature. Wood is a poor conductor. Heat applied to it is concentrated at the surface and only slowly penetrates to warm the interior. As mentioned earlier, different materials have differing heat capacities which is another surface property affecting temperature.

Moisture plays an important role in influencing surface temperatures. Water has a high specific heat capacity and is a good conductor. These properties make the moisture content of a material an important factor influencing its temperature. Compared to dry surfaces, moist surfaces will not reach as high of a maximum temperature, nor will they cool to as low of a minimum temperature. Therefore, moisture acts to moderate surface temperatures by reducing the daily temperature range. Winds also have a moderating effect. During the day, winds help transfer heat away from the surface, reducing maximum air temperatures. While at night, winds mix warmer air down to the surface thus raising minimum air temperatures. Vegetation also influences surface temperatures- usually moderating air temperatures within the vegetative layer (see: Effects of forests on surface temperatures).

The above discussion of surface properties focuses on factors that contribute to horizontal variations in surface air temperature; however, air temperature also varies with height. The processes that control the variation of surface temperatures with height are discussed in Lapse Rates and Inversions.

In all situations, vegetation moderates air temperatures within the vegetative layer for several reasons. First, it intercepts both incoming and outgoing radiation and therefore has a marked effect on ground temperature; second, green foliage does not warm up as much as ground or dry litter; and third, leaf surfaces exchange heat with air through a deeper, less restricted boundary layer. These effects result in less pronounced temperature changes with height above the ground.

In all vegetative cover, the temperature distribution depends upon the nature and density of the vegetation. With plants, such as low brush, the leaves form a nearly continuous upper surface, and this surface acts as the effective ground surface. The maximum daytime temperatures and minimum nighttime temperatures are near the top of the brush or dense plant cover, although temperatures near the ground are not greatly different.

The crowns of trees in a heavy forest form a nearly continuous cover and the canopy thus becomes, in effect, the air contact surface. The highest daytime temperatures are found near the crown top, and the temperature will decrease gradually between this level and the ground. Maximum air temperatures near the crowns may be 18° to 20° warmer than air temperature near the ground. Above the tree crowns the temperature decreases fairly rapidly with height, although never as rapidly as over bare ground. This is because the temperatures of the tree crown surfaces in contact with the air are lower than bare ground, and because the air circulation around these surfaces is better.

Less dense vegetation will permit more solar radiation to penetrate to the ground than will a dense cover. The degree of partial ground shading provided by less dense vegetation determines the air temperature distribution between the ground surface and the canopy top. It will range between that found over bare ground and that under a closed canopy.

Air temperatures at the standard 4 1/2-foot height within the forest in the afternoon are likely to be 5° to 8° cooler than the temperatures in nearby cleared areas. Openings in a moderate to dense timber stand may become warm air pockets during the day. These openings often act as natural chimneys and may accelerate the rate of burning of surface fires that are close enough to be influenced by these "chimneys."

Night temperatures in dense timber stands tend to be lowest near the top of the crown where the principal radiation takes place. Some cool air from the crowns sinks down to the ground surface, and there is some additional cooling at the surface by radiation to the cooling crowns. Sparse timber or other vegetation will merely decrease the strength of the inversion just above the ground surface.

Seasonal temperature patterns are affected principally by latitude, large water bodies, and the general circulation patterns. The latitude effect is due to the angle at which the suns rays strike the earth (see Radiation). In general, the seasonal variation of temperatures near the surface is least in equatorial regions, where there is little difference in solar heating through the year. This seasonal variation increases with latitude to both polar regions, where summer days have a maximum of 24 hours of sunshine and winter days a maximum of 24 hours of darkness. Large water bodies moderate the seasonal temperature cycle because of their great heat capacity. In one area, the general circulation pattern may produce cloudy weather with successive influxes of cold air, and thus a reduction in the monthly or seasonal temperature. In another area, the same pattern may produce opposite effects.

The diurnal temperature variation depends upon all of the factors we have discussed so far. The normal daily pattern at an inland location with level terrain consists of a daily temperature range of 20-30° F near the surface, with the highest temperature in midafternoon and the lowest temperature just after sunrise. This diurnal temperature range decreases with altitude above the surface. Various factors alter this pattern.

• A primary factor is the character of the surface. In general, those surfaces that become warmest during the day also become coldest at night, and the air temperature above them also has a high daily range. Snow surfaces are an important exception.

• Clouds, strong winds, high humidity, and atmospheric instability lower the maximum temperature and raise the minimum temperature, thereby reducing the daily temperature range.

• In mountainous terrain, one finds a greater diurnal variation in temperature in the valleys, and less along the slopes (in the thermal belt) and at higher elevations. Aspect affects the solar radiation and therefore the diurnal temperature range and the time of maxima and minima. Maxima will occur earlier, for example, on east slopes than on west slopes. Even minor shape characteristics of topography have their effects. Concave areas will have a larger daily range than convex areas.

• Large water bodies tend to moderate the daily temperature variation just as they moderate the seasonal variation. Coastal areas have a marine, rather than continental, climate.

Diurnal changes in temperature take place within the limitations of air-mass temperature. The passage of a front, evidence that another air mass has moved into the area, is reflected in the temperature pattern. Temperatures drop when a cold air mass moves in, and rise when a warm air mass moves in. In some cases the diurnal pattern is completely obscured. The temperature may continue to fall throughout the day when a very cold air mass moves in rapidly, or may continue to rise throughout the night when a warm air mass moves in. Along the west coast during the summer, a cool, marine air mass is usually found at low levels, and a warm, dry air mass is usually found above. A change in the vertical height of the boundary layer between these two air masses will appear in the temperature patterns along the slopes.

The atmosphere is primarily heated and cooled at the earths surface through conduction and convection and at higher levels in the atmosphere heat is lost to space through radiation. Therefore, we can expect temperatures to decrease with height. Another reason for a decrease in temperature with increasing height is that pressure decreases with height and decreases in pressure lead to decreases in temperature (see Temperature, Heat and the Atmosphere).

The average rate of decrease of temperature, or lapse rate, in the troposphere (approximately the lowest 7 miles of the atmosphere) is 3.5º F per 1000 ft or 6.3º C per kilometer. If solely the change in pressure were responsible for the decrease in temperature with altitude the lapse rate would be larger ( 5.5º F per 1000 ft or 9.8º C per kilometer ) and result in a more rapid drop in temperature with height; this is the adiabatic lapse rate. Lapse rates in the atmosphere can often deviate widely from the average value as different layers of air can have different temperatures as a result of horizontal motion aloft. Near the surface during the day, air at the ground can be substantially warmer than that just a few feet above the surface resulting in lapse rates far larger than the either the average or adiabatic lapse rate. This is often termed a superadiabatic lapse rate.

An inversion is a layer of the atmosphere where the air temperature increases with increasing height. At night when the earths surface is cooling the air in contact with it, the air temperature will actually increase with height over a shallow layer forming a night inversion. Marine inversions are common in coastal areas. Lapse rates and inversions are key components in determining atmospheric stability.

The above discussion focuses on how air temperature varies with height, however there are also horizontal variations in surface air temperature.

night inversion

smoke and inversions

Air cooled at night, primarily by contact with cold, radiating surfaces, gradually deepens as the night progresses and forms a surface inversion. This is a surface layer in which the temperature increases with height. Such an inversion may involve a temperature change of as much as 25° F. in 250 vertical feet. The cold air is dense and readily flows down slopes and gathers in pockets and valleys. Surface inversions forming at night are commonly referred to as night inversions. Night inversions are so important in fire behavior that we should consider them in some detail.

Night inversions are common during clear, calm, settled weather. They are usually easy to identify. Inversions trap impurities, smoke, and factory and traffic fumes, resulting in poor visibility. Smoke from chimneys rises until its temperature matches that of the surrounding air. Then it flattens out and spreads horizontally. If fog forms in the cold air, it is generally shallow ground fog, its depth generally less than that of the inversion. Ground fog in patches in surface depressions along highways is formed in small-scale inversions.

Cloudiness and water vapor in the atmosphere limit the formation and strength of night inversions by reducing the rate of outgoing radiation from the earth.

On windy nights, compared with calm nights, turbulence and mixing distribute the cooling through a deeper layer, and the temperature decrease is less. Winds may reduce and sometimes prevent the formation of a night inversion. The drop in temperature near the ground at night is thereby often abruptly stopped or even reversed when the wind picks up.

Night inversions are shallow but more intense when the overall temperature structure of the atmosphere is stable. Under unstable conditions, convection distributes available heat. Inversions are therefore less likely, and those occurring will not be as intense. However, if a night inversion is able to form, mixing is reduced in the lower layers.

Topography plays a decided role in both the formation and intensity of night inversions. Cold air layers are quite shallow on slopes and in open canyons or ravines where the cold, dense air can drain away as it is formed. This descent of cold air results in the formation of deep, cold layers and inversions in valleys. Inversion layers are both more common and intense in lower mountain valleys or in basins with poor air drainage, than in flat areas.

In mountainous areas, the height of the top of night inversions, although it varies from night to night, is usually below the main ridges. The height of the warmest air temperature at the inversion top can be found by measuring temperatures along the slope. From this level, the temperatures decrease as one goes farther up or down the slope. At this level are both the highest minimum temperatures and the least daily temperature variation of any level along the slope. Here also are the lowest nighttime relative humidity and the lowest nighttime fuel moisture. Because of these characteristics of the average level of the inversion top, it is known as the thermal belt. Within the thermal belt, wildfires can remain quite active during the night. Below the thermal belt, fires are in cool, humid, and stable air, often with downslope winds. Above the thermal belt, temperatures decrease with height. The effect of the lower temperatures, however, may be offset by stronger winds and less stable air as fires penetrate the region above the thermal belt.

Night inversions in mountainous country increase in depth during the night. They form early in the evening at the canyon bottom or valley floor and at first are quite shallow. Then the cold layer gradually deepens, the top reaching farther up the slope with the continued cooling from the surface and the flow of cold air from adjoining slopes. A maximum depth is reached during the middle of the night, and the depth may then remain constant or decrease slightly just before sunrise. If the air is sufficiently cold and moist, fog may form.

After sunrise, surface heating begins to warm the cold air, and the inversion top may actually rise slightly from this expansion. As heating destroys the inversion along the slopes, upslope winds begin. The transport of air from the valley bottom up the slopes may actually cause the inversion top to lower over the middle of the valley. Finally, with continued heating and mixing, the inversion layer is completely dissipated. The behavior of a fire burning beneath an inversion may change abruptly when the inversion is destroyed.

See also: Marine inversions.

marine inversion

A common type of warm-season inversion, found particularly along the west coast, is the coastal or marine inversion. Here cool, moist air from the ocean spreads over low-lying land. The layer of cool, moist air may vary in depth from a few hundred to several thousand feet. This layer is topped by a much warmer, drier, and relatively unstable air mass. Marine inversions, although they may persist in some areas during the day, are strongest and most noticeable at night. Fog and stratus clouds often form in the cool marine air at night and move inland into coastal basins and valleys. If the cold air is quite shallow, fog usually forms. If the layer is deep, stratus clouds are likely to form.

See also: Night inversions.

Before or during prescribed fires, the following sources of weather information can be used to forecast safe burning conditions. Ordinarily, four sources of weather information are available:

• National Weather Service
• State Forestry Agencies
• US Forest Service
• Local Observations
• Private Forecasting Services

National Weather Service

Local National Weather Service offices furnish weather forecasts and outlooks via radio and television. Spot weather forecasts are also available, but their value depends upon the forecasters knowledge of local conditions. Inexpensive radios are also available that continually monitor National Oceanic and Atmospheric Administration (NOAA) weather-related information and forecast updates. Relying solely on the NOAA broadcasts is ill-advised because this information is not specific enough for smoke-management planning.

State Forestry Agencies

The best source of information including current forecasts and outlooks is generally the local office of your state forestry agency. The person you talk to can often help you interpret the forecast, give you any warnings, and pass on pertinent information such as other burns planned for that day. The prescribed burner should take full advantage of such services.

All southern state forestry agencies and national forests, as well as many military bases and private concerns operate fire-danger stations. The basic weather parameters measured at these sites are very useful. However, National Fire Danger Rating System (NFDRS) indices which are calculated from these measurements should not be used as a planning tool for prescribed burning. This system was designed to provide a worst-case scenario for wildfire control over very large areas.

Local Observations

Weather observations should be made at the prescribed burn site immediately before, during, and immediately after a fire. Such observations are important because they serve as a check on the applicability of the forecast and keep the burning crew up-to-date on any local influences or changes. Take readings in a similar area upwind of the fire to avoid heating and drying effects from the fire. Do this at 1- to 2- hour intervals, or more often if changes in fire behavior are noticed. Measurements taken in an open area, on a forest road, and in a stand are likely to differ widely. Easy-to-use belt weather kits that include a psychrometer and an anemometer are available. By using this kit and observing cloud conditions, a competent observer can obtain a fairly complete picture of the current weather.

Moisture escapes into the atmosphere through evaporation from water bodies and soil, and through transpiration from vegetation. The dew-point temperature and the absolute humidity represent the actual moisture in the air, while the relative humidity indicates the degree of saturation at a given temperature. Absolute humidity varies in space and time for several reasons; however, relative humidity does not necessarily change in the same manner, because relative humidity is very dependent upon air temperature. The temperature effect frequently overrides the absolute humidity effect; therefore, relative humidity usually varies inversely with temperature.

Atmospheric moisture is a key element in fire weather. It has direct effects on the flammability of forest fuels, and, by its relationship to other weather factors, it has indirect effects on other aspects of fire behavior. There is a continuous exchange of water vapor between the atmosphere and dead wildland fuels. Dry fuels absorb moisture from a humid atmosphere and give up their moisture to dry air. During very dry periods, low humidity may also affect the moisture content of green fuels. When atmospheric moisture condenses and falls as precipitation, it increases the moisture content of dead fuels, and, by replenishing soil moisture, it provides for the growth of green vegetation.

Moisture influences surface temperatures, including surface fuel temperatures, by controlling radiation in its vapor state and by reflecting and radiating when it is condensed into clouds. The heat energy released in condensation provides the energy for thunderstorms and the violent winds associated with them. Moisture is also necessary for the development of lightning, which in many mountainous areas is a cause of wildfire.

• Sources of Atmospheric Moisture
• Evaporation
• Dew Point
• Absolute Humidity
• Relative Humidity
  • Variations in Relative Humidity Near the Ground
  • Variations in Relative Humidity With Height
• Measuring Humidity

Moisture as vapor acts the same as any other gas. It mixes with other gases in the air, and yet maintains its own identity and characteristics. It is the raw material in condensation. It stores immense quantities of energy gained in evaporation; this energy is later released in condensation. Much of the energy for thunderstorms, tornadoes, hurricanes, and other strong winds comes from the heat released when water vapor condenses. The availability of water vapor for precipitation largely determines the ability of a region to grow vegetation, which later becomes the fuel for wildland fires.

Moisture in the atmosphere is continually changing its physical state--condensing into liquid, freezing into ice, melting into liquid water, evaporating into gaseous water vapor, and condensing back to liquid. These changes are all related to temperature, the gage of molecular activity in any substance. At about -460° F (absolute zero) the molecules of all substances are motionless. As the temperature rises, they move around at increasing speeds. Water molecules move slowly at subfreezing temperatures, more rapidly at melting temperature, and still more rapidly through the boiling stage. However, at any given temperature, individual molecules, whether solid, liquid, or gas, do not have the same speeds or direction of travel. Collisions that change their speeds and directions occur continuously.

Evaporation

Some molecules momentarily acquire a very high speed from the impacts of other molecules. If this collision occurs in liquid water near the surface, and the high speed is in an outward direction, the molecules may escape into the air. This is evaporation, the process by which a liquid water molecule becomes a water-vapor molecule. Since molecules with the highest energy content escape, leaving behind in the liquid those with a lower energy content, the average level of energy of this liquid is decreased. The decrease in energy level results in a decrease in temperature of the liquid. Therefore, evaporation is a cooling process. Each molecule escaping into the air by a change of state takes with it nearly 1,000 times the energy needed to raise the temperature of a water molecule 1° F.

Vapor pressure

The pressure at the water-air boundary resulting from molecular motion in the direction of escape from the liquid is called the vapor pressure of water. This pressure varies only with the temperature of the water and determines the rate at which water molecules escape to the air and become vapor molecules. The water-vapor molecules which escape to the air displace air molecules and contribute their proportionate share to the total atmospheric pressure. This portion is called the partial pressure due to water vapor, or for simplicity, the vapor pressure.

Vapor pressure depends on the actual water vapor in the air, and it may vary from near zero in cold, dry air to about 2 inches of mercury in warm, moist air. High values can occur only in the warm, lower layers of the troposphere. The pressure produced by the vapor causes some water-vapor molecules to re-enter water surfaces by condensation. The same amount of heat energy that was needed for evaporation is liberated to warm the condensation surface.

At the water-air boundary, molecules are exchanged in both directions continuously, but the exchange is usually greater in one direction or the other. Evaporation occurs when more molecules leave the water surface than enter it, and condensation occurs when the opposite takes place. Actually, both condensation and evaporation occur at the same time. As noted earlier, a similar exchange of molecules takes place between water vapor and ice in the process of sublimation. The vapor pressure of ice is somewhat less than that of water at the same temperature. Hence, at low temperatures sublimation on ice is accomplished more readily than condensation on a water surface.

When the vapor pressure in the atmosphere is in equilibrium with the vapor pressure of a water or ice surface, there is no net exchange of water molecules in either direction, and the atmosphere is said to be saturated. A saturated volume of air contains all the vapor that it can hold. The vapor pressure at saturation is called the saturation vapor pressure. The saturation vapor pressure varies with the temperature of the air and is identical to the vapor pressure of water at that temperature. The higher the temperature, the more water vapor a volume of air can hold, and the higher the saturation vapor pressure. Conversely, the lower the temperature, the lower the saturation vapor pressure. Table: Saturation Vapor Pressure illustrates how the saturation vapor pressure varies with temperature. In the common range of temperatures in the lower atmosphere, the saturation vapor pressure just about doubles for each 20^o F increase in temperature. With this understanding of evaporation, condensation, and vapor pressure, we can now define several terms used to indicate the amount of moisture in the atmosphere.

The air near the surface is usually not saturated; therefore, the actual vapor pressure is usually less than the saturation vapor pressure. The actual vapor pressure can be raised to saturation vapor pressure by evaporating more moisture into the air, or, since saturation vapor pressure varies with temperature, the air can be cooled until the saturation vapor pressure is equal to the actual vapor pressure. Evaporation alone does not ordinarily saturate the air except very close to the evaporating surface. Normal circulation usually carries evaporated moisture away from the evaporating surface.

A saturated volume of air contains all the water vapor that it can hold. Saturation is usually reached by the air being cooled until its saturation vapor pressure equals the actual vapor pressure. The temperature of the air at that point is called the dew-point temperature, or simply, the dew point. If the air is cooled below its dew point, condensation occurs because the amount of water vapor in the air exceeds the maximum amount that can be contained at the lower temperature. Further cooling therefore causes clouds, fog, or dew to form. Under ordinary circumstances the actual vapor pressure cannot exceed the saturation vapor pressure by more than a very small amount.

Cooling near the surface normally results from contact with cool ground or water. Cooling to the dew point may also occur by lifting moist air to higher altitudes; it is thus cooled adiabatically. For example, consider air with a temperature of 80°F. and a vapor pressure of 0.362 inches of mercury. Referring to Table: Saturation Vapor Pressure, we find that if the air is cooled to 50°, the actual vapor pressure will equal the saturation vapor pressure. Therefore, 50° is the dew point.

The actual amount of water vapor in a given volume of air, that is, the weight per volume, such as pounds per 1,000 cubic feet, is called the absolute humidity. A direct relationship exists among the dew point, the vapor pressure, and the absolute humidity because, at constant atmospheric pressure, each of these depends only on the actual amount of water vapor in the air. At saturation, the dew point is the same as the temperature, the vapor pressure is the saturation vapor pressure, and the absolute humidity is the saturation absolute humidity.

Table:Dew Point, Vapor Pressure, and Absolute Humidity shows the relationship among these three measures of atmospheric moisture. Saturation values of vapor pressure and absolute humidity can be obtained by entering temperature instead of dew point in the first column. Because of these relationships, the temperature of the dew point is a convenient unit of measure for moisture. Air temperature and dew point accurately define atmospheric moisture at any time or place.

Absolute humidity varies in space and time for several reasons; however, relative humidity does not necessarily change in the same manner, because relative humidity is very dependent upon air temperature. The temperature effect frequently overrides the absolute humidity effect; therefore, relative humidity usually varies inversely with temperature.

Variations in Absolute Humidity

The actual amount of moisture in air will vary from one air mass to another, and even within an air mass there will be continuing variations in time and space.

The moisture contents of air masses are basically related to their regions of origin. Air masses originating in continental areas are relatively dry. Those coming from large bodies of water, such as the Atlantic or the Gulf of Mexico are moist. Those from the Pacific are moist or moderately moist. As these maritime air masses invade the continent, land stations will observe abrupt rises in absolute humidity. As any air mass traverses areas different from its source region, gradual changes take place as evaporation, transpiration, condensation, and precipitation add or subtract moisture.

Through a deep layer within an air mass, the absolute humidity, like the temperature, usually decreases with height. There are several reasons for this distribution. First, moisture is added to the atmosphere from the surface and is carried upward by convection and upslope and up valley winds. Second, when air is lifted, the water vapor, as well as the air, expands proportionately so that the moisture in any given volume becomes less and less. Thus, the absolute humidity decreases as the air is lifted. Third, since temperature usually decreases upward, the capacity for air to hold moisture decreases upward. Finally, the precipitation process removes condensed moisture from higher levels in the atmosphere and deposits it at the surface.

The normal pattern of decrease of moisture with altitude may be altered occasionally when horizontal flow at intermediate levels aloft brings in moist air. Such flow is responsible for much of the summer thunderstorm activity over large parts of the West. Extremely low absolute humidity is found in subsiding air aloft. This dry air originates near the top of the troposphere and slowly sinks to lower levels. If it reaches the ground, or is mixed downward, it may produce acutely low humidity near the surface and an abrupt increase in fire danger (See: Subsidence).

If we consider only a very shallow layer of air near the surface, we find that the vertical variation of absolute humidity with height will change during each 24-hour period as conditions favoring evaporation alternate with conditions favoring condensation. During clear days, moisture usually is added to the air by evaporation from warm surfaces; therefore, the absolute humidity decreases upward.

At night, moisture is usually taken from the air near the surface by condensation on cold surfaces and absorption by cold soil and other substances; thus, the absolute humidity may increase upward through a very shallow layer.

Saturation of surface air is a condition of favorable fire weather; that is, conducive to low fire danger. Less favorable are conditions of unsaturation, which permit evaporation from forest fuels, increasing their flammability and the fire danger. Therefore, a very useful measure of atmospheric moisture is the relative humidity. It is the ratio, in percent, of the amount of moisture in a volume of air to the total amount which that volume can hold at the given temperature and atmospheric pressure. Relative humidity is also the ratio of actual vapor pressure to saturation vapor pressure, times 100. It ranges from 100 percent at saturation to near zero for very dry air. Relative humidity depends on the actual moisture content of the air, the temperature, and the pressure.

The dependence of relative humidity on temperature must be kept in mind. Suppose that we have air at 80°F. and 24 percent relative humidity. Using Table:Dew Point, Vapor Pressure, and Absolute Humidity, we find that the saturation vapor pressure for 80° is 1.032 inches of mercury. We can compute the actual vapor pressure by multiplying 1.082 by 0.24. The actual vapor pressure is 0.248 rounded off. The dew point for this vapor pressure is 40°. We now know that if the air was cooled from 80°F. to 40°, with no other change, the humidity would increase from 24 percent to 100 percent and the air would be saturated. At that temperature the actual vapor pressure would equal the saturation vapor pressure. The absolute humidity in the table could be used in a similar manner. Thus; the relative humidity may change considerably with no addition of moisture, just by cooling alone.

Factors influencing variations in relative humidity

Relative humidity is much more variable than absolute humidity. It often changes rapidly and in significant amounts from one hour to the next and from place to place. Relative humidity is much more variable because it depends not only on absolute humidity but also on air temperature. It varies directly with moisture content and inversely with temperature. Because of these relationships, it is often not possible to make general statements about relative humidity variations, particularly vertical variations within short distances above the ground.

• Variations in Relative Humidity Near the Ground
• Variations in Relative Humidity With Height

Relative humidity is most important as a fire-weather factor in the layer near the ground, where it influences both fuels and fire behavior. Near the ground, air moisture content, season, time of day, slope, aspect, elevation, clouds, and vegetation all cause important variations in relative humidity.

Diurnal changes in relative humidity

Since hourly and daily changes of relative humidity are normally measured in a standard instrument shelter, we will consider variations at that level and infer from our knowledge of surface temperatures what the conditions are near the surface around forest fuels.

A typical fair-weather pattern of relative humidity, as shown on a hygrothermograph exposed in a shelter at a valley station or one in flat terrain, is nearly a mirror image of the temperature

pattern. Maximum humidity generally occurs about daybreak, at the time of minimum temperature. After sunrise, humidity drops rapidly and reaches a minimum at about the time of maximum temperature. It rises more gradually from late afternoon through the night. The daily range of humidity is usually greatest when the daily range of temperature is greatest. Variations in the humidity traces within an air mass from one day to the next are usually small, reflecting mostly differences in temperatures. But over several days, there may be noticeable cumulative differences in humidity as the air mass gradually picks up or loses moisture.

Seasonal changes in relative humidity

Seasonal changes in relative humidity patterns are also apparent. In western fire-weather seasons that begin following a moist spring and continue through the summer and early fall, a seasonal change is particularly noticeable. Daily temperature ranges are greatest early in the fire season when the sun is nearly overhead and night skies are clear. Strong nighttime cooling, in combination with ample moisture in the soil and vegetation to contribute moisture to the atmosphere, often boosts night humidities to or near 100%. Intensive daytime surface heating and convective transport of moisture upward combine to drop the relative humidity to low levels in the afternoon.

As the season progresses, soil and vegetation dry out and solar heating diminishes as the sun tracks farther south. Daytime humidities become even lower late in the season, but, with a greater reduction in night humidities, the daily range is reduced, and the fire weather is further intensified. Occasional summer rains may interrupt this progression but do not greatly change the overall seasonal pattern.

In areas that have separate spring and fall fire seasons, the daily temperature extremes are generally not so striking. Also, the cumulative drying of soil and vegetation is not so consistent, except during unusual drought. Because periodic rains generally occur during the seasons, the humidity changes tend to be somewhat variable. In some areas, seasonal increases in relative humidity decrease fire danger during the summer. In the Great Lakes region, particularly, where the many small lakes become quite warm during the summer and transpiration from vegetation is at its peak, daytime relative humidities do not reach as low values in the same air mass types as they do in spring and fall.

Effects of topography on humidity

Relative Humidity and Upper Elevations

Humidity may vary considerably from one spot to another, depending greatly on the topography. In relatively flat to rolling terrain, the humidity measured at a well-exposed station may be quite representative of a fairly large area. There will be local exceptions along streams, irrigated fields, in shaded woods, or in barren areas. In the daytime particularly, circulation and mixing are usually sufficient to smooth out local effects over relatively short distances.

In mountainous topography, the effects of elevation and aspect become important, and humidities vary more than over gentle terrain. Low elevations warm up and dry out earlier in the spring than do high elevations. South slopes also are more advanced seasonally than north slopes. As the season progresses, cumulative drying tends to even out these differences since stored moisture in the surface is depleted, but the differences do not disappear.

The decrease with height of both temperature and dew point produces higher relative humidities at higher elevations on slopes. The pattern is complicated, however, because of heating of the air next to the slopes, the transport of moisture with upslope winds, and the frequent stratification of moisture into layers, so generalizations are difficult to make.

When nighttime cooling begins, the temperature change with height is usually reversed. Cold air flowing down the slopes accumulates at the bottom. As the night progresses, additional cooling occurs, and by morning, if the air becomes saturated, fog or dew forms. Relative humidity may decrease from 100% at the foot of the slope to a minimum value at the top of the temperature inversion and then may increase slightly farther up the slope above the inversion.

Just as south slopes dry out faster because of their higher day temperatures, they also have somewhat lower day relative humidities than north slopes throughout the summer. At upper elevations, though, the difference between north and south slopes becomes negligible because of the good air mixing at these more exposed sites. At night, humidity differences on north and south slopes become slight.

In most mountainous country, the daily range of relative humidity is greatest in valley bottoms and least at higher elevations. Thus, while fires on lower slopes may burn better during the day, they often quiet down considerably at night when humidity increases. But at higher elevations, particularly in and above the thermal belt, fires may continue to burn aggressively through the night as humidities remain low, temperatures stay higher, and wind speed is greater.

Again, we should be cautious of generalizations. For example, in the summer in the Pacific coast ranges, higher humidities are usually found on ridgetops during the day than during the night. This anomaly results from slope winds carrying moisture upward from the moist marine air layer during the day. Moist air that is not carried away aloft settles back down at night.

Effect of wind on humidity

Wind mixes evaporating water vapor with surrounding air and evens out temperature extremes by moving air away from hot and cold surfaces. Thus, diurnal ranges of relative humidity are less during windy periods than during calm periods. Winds also reduce place-to-place differences by mixing air of different moisture contents and different temperatures. Patches of fog on a calm night indicate poor ventilation.

Effect of clouds on humidity

Clouds strongly affect heating and cooling and therefore influence the relative humidity. The humidity will be higher on cloudy days and lower on cloudy nights. Thus, clouds reduce the daily range considerably. Precipitation in any form raises relative humidities by cooling the air and by supplying moisture for evaporation into the air.

Effect of vegetation on humidity

Vegetation moderates surface temperatures and contributes to air moisture through transpiration and evaporation--both factors that affect local relative humidity. A continuous forest canopy has the added effect of decreasing surface wind speeds and the mixing that takes place with air movement.

The differences in humidity between forest stands and open areas generally vary with the density of the crown canopy. Under a closed canopy, humidity is normally higher than outside during the day, and lower at night. The higher daytime humidities are even more pronounced when there is a green understory. Deciduous forests have only slight effects on humidity during their leafless period.

Two factors lessen the humidity difference between forest stands and forest openings. Overcast skies limit both heating and cooling, and drought conditions decrease the amount of moisture available for evaporation and transpiration.

Openings of up to about 20 yards in diameter do not have daytime relative humidities much different from under the canopy--except at the heated ground surface. These openings serve as chimneys for convective airflow, and surface air is drawn into them from the surrounding forest (see: Surface air temperatures). At night in small openings, the stagnation coupled with strong radiation can cause locally high humidities.

The daytime humidities in larger clearings are much like those in open country. If the airflow is restricted, however, temperatures may rise slightly above those at exposed stations, and humidities will be correspondingly lower. In the afternoon, these may range from 5 to 20% lower in the clearing than within a well-shaded forest. Night humidities are generally similar to those at exposed sites, usually somewhat higher than in the woods.

Open forest stands have humidity characteristics somewhere between those of exposed sites and closed stands, depending on crown density. During dry weather, especially after prolonged dry spells, the differences in relative humidity between forested and open lands become progressively less.

Effect of air masses on humidity

The amount of moisture in the air is one of the air-mass characteristics. Air masses originating over water bodies will have higher moisture contents than those originating over continents. When a front passes, and a different air mass arrives, a change in absolute humidity can be expected. The change in relative humidity, however, will depend greatly on the air-mass temperature. A warm, dry air mass replacing a cool, moist one, or vice versa, may cause a large change in relative humidity. A cool, dry air mass replacing a warm, moist one, however, may actually have a higher relative humidity if its temperature is appreciably lower.

Along the west coast, when a lower marine layer is topped by a warm, dry, subsiding air mass, the inversion layer is actually the boundary between two very different air masses. Inland, where the inversion intersects the coast ranges, very abnormal relative humidity patterns are found. In these inland areas, the inversion is usually higher in the day and lower at night; however, along the coastal low-lands, the reverse is usually true. Along the slopes of the adjacent mountains, some areas will be in the marine air during the day and in the dry, subsiding air at night. The relative humidity may begin to rise during the late afternoon and early evening and then suddenly drop to low values as dry air from aloft moves down the slopes. Abrupt humidity drops of up to 70% in the early evening have been observed.

Relative Humidity

During the day near the surface, particularly with clear skies, both the temperature and absolute humidity usually decrease with height. These two variables have opposite effects on the relative humidity. Which effect is dominant depends upon the dryness of the surface. The relative humidity usually increases with height over normal surfaces because the effect of the decrease in temperature is greater than that of the decrease in absolute humidity. Over a moist surface, however, the effect of the decrease in absolute humidity may overbalance that of temperature decrease, and the relative humidity in the surface layer will decrease with height.

At night, the change of temperature with height usually predominates, and the relative humidity will decrease with height through the lowest layers.

Above the lowest layers, the relative humidity generally increases with height in the day through much of the lower troposphere. Convection alone would account for this increase. As air is lifted, the temperature decreases 5.5° F per 1,000 feet, and the dew point decreases at about 1° F per 1,000 feet. Therefore, the dew point and the temperature become 4.5° F closer per 1,000 feet, and the relative humidity increases until saturation is reached.

A subsiding layer of air in the troposphere warms by the adiabatic process and forms a subsidence inversion. The relative humidity will decrease upward through the temperature inversion at the base of the subsiding layer. The marine inversion along the west coast, for example, is a subsidence inversion. The marine air below has low temperatures and high humidities, and the adiabatically heated subsiding air mass above has higher temperatures and lower humidities. This pronounced change in temperature and humidity is evident along the slopes of coastal mountains when the marine inversion is present.

The most widely used device for accurately measuring atmospheric moisture near the surface is the psychrometer. It consists of two identical mercurial thermometers. One thermometer is used for measuring the air temperature; the other measures the temperature of evaporating water contained in a muslin wicking surrounding the thermometer bulb. The amount that the evaporating surface will cool is determined by the difference between the vapor pressure and the saturation vapor pressure. The first reading is commonly referred to as the dry-bulb temperature and the second as the wet-bulb temperature. The wet-bulb temperature is the steady value reached during a period of brisk ventilation of the thermometer bulbs. If the air is saturated, the wet-bulb and dry-bulb temperatures are the same.

From the wet- and dry-bulb measurements, computed values of dew-point temperature, absolute humidity, and relative humidity may be read from tables or slide rules. As noted earlier, these moisture relations vary with changes in pressure. The daily pressure changes as shown by the barometer are not large enough to be important, but those due to differences in elevation are significant. They have been considered in the construction of the tables or slide rules. The ones labeled with the correct pressure must be used.

Other instruments used to measure relative humidity contain fibers of various materials that swell or shrink with changing relative humidity. One instrument of this type that records a continuous trace of relative humidity is called a hygrograph. A more common form in use at fire-weather stations is the hygrothermograph, which records both relative humidity and temperature. Other devices, such as those commonly used for upper-air soundings, employ moisture-sensitive elements that change in electrical or chemical characteristics with changing humidity.

Standard surface measurements of relative humidity, like those of temperature, are made in an instrument shelter 4-1/2 feet above the ground. A properly operated sling psychrometer, however, will indicate dry- and wet-bulb readings that agree well with those obtained in the shelter. The only necessary precautions are to select a well-ventilated shady spot, and to whirl the instrument rapidly for a sufficient time to get the true (lowest) wetbulb temperature. Care must be taken not to allow the wicking to dry out, and not to break the thermometer by striking any object while whirling the psychrometer.

Measuring relative humidity near the ground

The relative humidity that affects fuels on the forest floor is often quite different from that in the instrument shelter, particularly in unshaded areas where soil and surface fuels exposed to the sun are heated intensely, and warm the air surrounding them. This very warm air may have a dew point nearly the same or slightly higher than the air in the instrument shelter, but because it is much warmer, it has a much lower relative humidity.

With similar exposure at night, humidities are likely to be higher near the ground than in the shelter because of radiative cooling of the surface. Often, dew will form on the surface, indicating 100 percent relative humidity, when the humidity at shelter height may be considerably below the saturation level.

These conditions are typical for relatively still air, clear skies, and open exposure. When wind speeds reach about 8 miles per hour, the increased mixing diminishes the difference between surface and shelter-height humidities Also, under heavy cloud cover or shade, the humidity differences between the two levels tend to disappear because the principal radiating surface is above both levels.

It is impractical to measure humidity close to the ground with field instruments, but with the aid of tables, the humidity can be estimated from psychrometric readings at the standard height and a dry-bulb temperature reading at the surface. We must assume that the dew point is the same at both levels. Although we know that this may not be exact, it will give a reasonable estimation.

Consider the example for a pressure of 29 inches.

The 8-percent relative humidity was obtained from a complete set of tables, using a dry-bulb temperature of 140°F. and a dew point of 56°F.

Atmospheric stability is the resistance of the atmosphere to vertical motion. The distribution of temperature vertically in the troposphere influences vertical motion. A large decrease of temperature with height indicates an unstable condition which promotes up and down currents. A small decrease with height indicates a stable condition which inhibits vertical motion. Where the temperature increases with height, through an inversion, the atmosphere is extremely stable.

To determine stability conditions, temperature lapse rates are compared to dry- or moist-adiabatic lapse rates. Between stable and unstable lapse rates we may have a conditionally unstable saturated. During condensation in saturated air, heat is released which warms the air and may produce instability; during evaporation, heat is absorbed and may increase stability.

Several different lifting processes contribute to atmospheric stability, such as local heating, with wind speed, surface characteristics, warm- and cold-air advection, and many other factors. As a result, atmospheric stability varies with these factors. Atmospheric stability also varies diurnally and seasonally. We can use type of cloud, windflow characteristics, occurrence of dust devils, and other phenomena as indicators of stability.

Subsidence is the gradual lowering of a layer of air over a broad area. When it begins at high levels in the troposphere, the air, which has little initial moisture, becomes increasingly warmer with resulting lower relative humidity as it approaches the surface. If some mechanism is present by which this warm, dry air can reach the surface, a very serious fire situation can result.

Atmospheric stability is closely related to fire behavior, and a general understanding of stability and its effects is necessary to the successful interpretation of fire-behavior phenomena. Atmospheric stability may either encourage or suppress vertical air motion. The heat of fire itself generates vertical motion, at least near the surface, but the convective circulation thus established is affected directly by the stability of the air. In turn, the indraft into the fire at low levels is affected, and this has a marked effect on fire intensity. Also, in many indirect ways, atmospheric stability will affect fire behavior. For example, winds tend to be turbulent and gusty when the atmosphere is unstable, and this type of airflow causes fires to behave erratically. Thunderstorms with strong updrafts and downdrafts develop when the atmosphere is unstable and contains sufficient moisture. Their lightning may set wildfires, and their distinctive winds can have adverse effects on fire behavior.

• Determining Atmospheric Stability
• Lifting Processes
• Diurnal and Seasonal Variations in Stability
• Local Indicators of Stability
• Subsidence

The degree of stability or instability of an atmospheric layer is determined by comparing its temperature lapse rate, as shown by a sounding, with the appropriate dry- or moist-adiabatic lapse rates. The dry adiabatic rate is used for air that is not saturated and the moist-adiabatic rate is used for saturated air. The adiabatic process is reversible. Just as air expands and cools when it is lifted, so is it equally compressed and warmed as it is lowered. Hence, stability determinations for either upward or downward moving air parcels make use similar comparisons with the appropriate adiabatic lapse rates.

Unsaturated air

Determining Stability

• Stable: A temperature lapse rate less than the dry adiabatic rate of 5.5° F per 1,000 feet for an unsaturated parcel is considered stable, because vertical motion is damped.
• Unstable: A lapse rate greater than dry-adiabatic (a superadiabatic lapse rate) favors vertical motion and is unstable. Since it is unstable, the air tends to adjust itself through mixing and overturning to a more stable condition. Superadiabatic lapse rates are not ordinarily found in the atmosphere except near the surface of the earth on sunny days. When an unsaturated layer of air is mixed thoroughly, its lapse rate tends toward neutral stability.
• Neutral: In the absence of saturation, an atmospheric layer is neutrally stable if its lapse rate is the same as the dry-adiabatic rate. Under this particular condition, any existing vertical motion is neither damped nor accelerated.

Warming of the lower layers during the daytime by contact with the earths surface or by heat from a wildfire will make a neutral lapse rate become unstable. In an atmosphere with a dry-adiabatic lapse rate, hot gases rising from a fire will encounter little resistance, will travel upward with ease, and can develop a tall convection column. A neutrally stable atmosphere can be made unstable also by advection; that is, the horizontal movement of colder air into the area aloft or warmer air into the area near the surface. Once the lapse rate becomes unstable, vertical currents are easily initiated. Advection of warm air aloft or cold air near the surface has the reverse effect of making the atmosphere more stable.

The term "neutral" stability sounds rather passive, but we should be cautious when such a lapse rate is present. The temperature structure of the atmosphere is not static, but is continually changing. Any warming of the lower portion or cooling of the upper portion of a neutrally stable layer will cause the layer to become unstable, and it will then not only permit, but will assist, vertical motion. Such changes are easily brought about. Thus, we should consider the terms stable, neutral, and unstable in a relative, rather than an absolute, sense. A stable lapse rate that approaches the dry-adiabatic rate should be considered relatively unstable.

Saturated air

Adiabatic Chart

So far we have considered adiabatic cooling and warming and the degree of stability of the atmosphere only with respect to air that is not saturated. Rising air, cooling at the dry-adiabatic lapse rate, may eventually reach the dew-point temperature. Further cooling results in the condensation of water vapor into clouds, a change of state process that liberates the latent heat contained in the vapor. This heat is added to the rising air, with the result that the temperature no longer decreases at the dry-adiabatic rate, but at a lesser rate which is called the moist-adiabatic rate. On the average, this rate is around 3° F per 1,000 feet, but it varies slightly with pressure and considerably with temperature. The variation of the rate due to temperature may range from about 20° F per 1,000 feet at very warm temperatures to about 5° F per 1,000 feet at very cold temperatures. In warmer air masses, more water vapor is available for condensation and therefore more heat is released, while in colder air masses, little water vapor is available.

To determine the degree of stability or instability for a saturated parcel, the same stability terms apply as for unsaturated but the comparison of atmospheric lapse rate is made with the moist-adiabatic rate appropriate to the temperature encountered.

• Stable: A temperature lapse rate less than the moist adiabatic rate is considered stable.
• Unstable: A lapse rate greater than moist-adiabatic is unstable.
• Neutral: In the absence of saturation, an atmospheric layer is neutrally stable if its lapse rate is the same as the dry-adiabatic rate. Under this particular condition, any existing vertical motion is neither damped nor accelerated.

Conditional stability

An atmosphere that has a lapse rate lying between the dry and moist adiabats is said to be conditionally unstable. It is stable with respect to a lifted air parcel as long as the parcel remains unsaturated, but it is unstable with respect to a lifted parcel that has become saturated.

A saturated parcel in free convection loses additional moisture by condensation as it rises. This, plus the colder temperature aloft, causes the moist-adiabatic lapse rate to increase toward the dry-adiabatic rate. The rising parcel will thus eventually cool to the temperature of the surrounding air where the free convection will cease. This may be in the vicinity of the tropopause or at some lower level, depending on the temperature structure of the air aloft.

Layer Stability

Atmospheric stability is affected by vertical movement of both parcels of air or whole layers of considerable horizontal extent of the atmosphere. When an entire layer of stable air is lifted it becomes increasingly less stable. The layer stretches vertically as it is lifted, with the top rising farther and cooling more than the bottom. If no part of the layer reaches condensation, the stable layer will eventually become dry-adiabatic. Occasionally, the bottom of a layer of air being lifted is moister than the top and reaches its condensation level early in the lifting. Cooling of the bottom takes place at the slower moist-adiabatic rate, while the top continues to cool at the dry-adiabatic rate. The layer then becomes increasingly less stable at a rate faster than if condensation had not taken place. A descending (subsiding) layer of stable air becomes more stable as it lowers. The layer compresses, with the top sinking more and warming more than the bottom. The adiabatic processes involved are just the opposite of those that apply to rising air.

Several processes can cause air to be lifted into the atmosphere: convection, orographic lifting, frontal lifting, turbulence, and convergence.

Convection

A common process by which air is lifted in the atmosphere is convection. If the atmosphere remains stable, convection will be suppressed. But surface heating makes the lower layers of the atmosphere unstable during the daytime. Triggering mechanisms are required to begin convective action, and they usually are present. If the unstable layer is deep enough, so that the rising parcels reach their condensation level, cumulus-type clouds will form and may produce showers or thunderstorms if the atmospheric layer above the condensation level is conditionally unstable. Wildfire also may be a source of heat, which will initiate convection. At times, the fire convection column will reach the condensation level and produce clouds. Showers, though rare, have been known to occur.

Orographic lifting

Orographic Lifting

Layers of air commonly flow in response to pressure gradients. In doing so, if they are lifted up and over mountains, they are subjected to what is called orographic lifting. This is a very important process along our north-south mountain ranges in the western regions and the Appalachians in the East, because the general airflow is normally from a westerly direction. If the air is initially stable, and if no condensation takes place, it sinks back to its original level after passing over a ridge. If it is neutrally stable, the air will remain at its new level after crossing the ridge. In an unstable atmosphere, air given an initial uplift in this way keeps on rising, seeking a like temperature level, and is replaced by sinking colder air from above. If the condensation level is reached in the lifting process, and clouds form, initially stable air can become unstable. In each case, the internal depth and lapse rate of the layer will respond as indicated above.

Frontal lifting

Warmer, lighter air layers frequently flow up and over colder, heavier air masses. This is referred to as frontal lifting and is similar in effect to orographic lifting (see Air masses and fronts). Stable and unstable air masses react the same way regardless of whether they are lifted by the slope of topography or by the slope of a heavier air mass.

Turbulence

Turbulence associated with strong winds results in mixing of the air through the turbulent layer. In this process, some of the air near the top of the layer is mixed downward, and that near the bottom is mixed upward, resulting in an adiabatic layer topped by an inversion. At times, the resultant cooling near the top of the layer is sufficient to produce condensation and the formation of stratus, or layer-like, clouds.

Convergence and divergence

The airflow around surface low-pressure areas in the Northern Hemisphere is counterclockwise and spirals inward. This inflow produces upward motion in low-pressure areas. Airflow into a Low from all sides is called convergence. Now, the air must move. It is prevented from going downward by the earths surface, so it can only go upward. Thus, low-pressure areas on a surface weather map are regions of upward motion in the lower atmosphere.

In surface high-pressure areas, the airflow is clockwise and spirals outward. This airflow away from a High is called divergence. The air must be replaced, and the only source is from aloft. Thus, surface high-pressure areas are regions of sinking air motion from aloft, or subsidence.

Frequently, two or more of the above processes will act together. For example, the stronger heating of air over ridges during the daytime, compared to the warming of air at the same altitude away from the ridges, can aid orographic lifting in the development of deep convective currents, and frequently cumulus clouds, over ridges and mountain peaks. Similarly, orographic and frontal lifting may act together, and frontal lifting may combine with convergence around a Low to produce more effective upward motion.

Stability frequently varies through a wide range in different layers of the atmosphere for various reasons. Layering aloft may be due to an air mass of certain source-region characteristics moving above or below another air mass with a different temperature structure. The inflow of warmer (less dense) air at the bottom, or colder (more dense) air at the top of an air mass promotes instability, while the inflow of warmer air at the top or colder air at the surface has a stabilizing effect. At lower levels, stability of the air changes with surface heating and cooling, amount of cloud cover, and surface wind all acting together. We will consider first the changes in stability that take place during a daily cycle and the effects of various factors; then we will consider seasonal variations.

Diurnal variations in stability

Diurnal changes in surface heating and cooling produce daily changes in stability, from night inversions to daytime superadiabatic lapse rates, that are common over local land surfaces. During a typical light-wind, fair-weather period, radiation cooling at night forms a stable inversion near the surface, which deepens until it reaches its maximum development at about daybreak. After sunrise, the earth and air near the surface begin to heat, and a shallow superadiabatic layer is formed. Convective currents and mixing generated in this layer extend up to the barrier created by the inversion. As the day progresses, the unstable superadiabatic layer deepens, and heated air mixing upward creates an adiabatic layer, which eventually eliminates the inversion completely. This usually occurs by mid or late morning. Active mixing in warm seasons often extends the adiabatic layer to 4,000 or 5,000 feet above the surface by midafternoon. The superadiabatic layer, maintained by intense heating, is usually confined to the lowest few hundreds of feet, occasionally reaching 1,000 to 2,000 feet over bare ground in midsummer.

As the sun sets, the ground cools rapidly under clear skies and soon a shallow inversion is formed. The inversion continues to grow from the surface upward throughout the night as surface temperatures fall. The air within the inversion becomes increasingly stable. Vertical motion in the inversion layer is suppressed, though mixing may well continue in the air above the inversion. This mixing allows radiational cooling above the inversion to lower temperatures in that layer only slightly during the night.

This diurnal pattern of nighttime inversions and daytime superadiabatic layers near the surface can be expected to vary considerably. Clear skies and low air moisture permit more intense heating at the surface by day and more intense cooling by radiation at night than do cloudy skies. The lower atmosphere tends to be more unstable on clear days and more stable on clear nights.

Strong winds diminish or eliminate diurnal variations in stability near the surface. Turbulence associated with strong wind results in mixing, which tends to produce a dry-adiabatic lapse rate. Mechanical turbulence at night prevents the formation of surface inversions, but it may produce an inversion at the top of the mixed layer. During the day, thermal turbulence adds to the mechanical turbulence to produce effective mixing through a relatively deep layer. Consequently, great instability during the day, and stability at night occur when surface winds are light or absent.

Stability in the lower atmosphere varies locally between surfaces that heat and cool at different rates. Thus, dark-colored, barren, and rocky soils that reach high daytime temperatures contribute to strong daytime instability and, conversely, to strong stability at night. Areas recently blackened by fire are subject to about the maximum diurnal variation in surface temperature and the resulting changes in air stability. Vegetated areas that are interspersed with openings, outcrops, or other good absorbers and radiators have very spotty daytime stability conditions above them.

Topography also affects diurnal changes in the stability of the lower atmosphere. Air in mountain valleys and basins heats up faster during the daytime and cools more rapidly at night than the air over adjacent plains. This is due in part to the larger area of surface contact, and in part to differences in circulation systems in flat and mountainous topography. The amount of air heating depends on orientation, inclination, and shape of topography, and on the type and distribution of ground cover. South-facing slopes reach higher temperatures and have greater instability above them during the day than do corresponding north slopes. Both cool about the same at night (see Slope and Valley Winds).

Instability resulting from superheating near the surface is the origin of many of the important convective winds. On mountain slopes, the onset of daytime heating initiates upslope wind systems. The rising heated air flows up the slopes and is swept aloft above the ridgetops in a more-or-less steady stream.

Over level ground, heated surface air, in the absence of strong winds to disperse it, can remain in a layer next to the ground until it is disturbed. The rising air frequently spirals upward in the form of a whirlwind or dust devil. In other cases, it moves upward as intermittent bubbles or in more-or-less continuous columns. Pools of superheated air may also build up and intensify in poorly ventilated valleys to produce a highly unstable situation. They persist until released by some triggering mechanism which overcomes inertia, and they may move out violently.

Seasonal variations in stability

The amount of solar radiation received at the surface during the summer is considerably greater than in the winter. This is due to the difference in solar angle and the duration of sunshine. Temperature profiles and stability reflect seasonal variation accordingly. In the colder months, inversions become more pronounced and more persistent, and superadiabatic lapse rates occur only occasionally. In the summer months, superadiabatic conditions are the rule on sunny days. Greater variation in stability from day to day may be expected in the colder months because of the greater variety of air masses and weather situations that occur during this stormy season.

In addition to the seasonal effects directly caused by changes in solar radiation, there is also an important effect that is caused by the lag in heating and cooling of the atmosphere as a whole. The result is a predominance of cool air over warming land in the spring, and warm air over cooling surfaces in the fall. Thus, the steepest lapse rates frequently occur during the spring, whereas the strongest inversions occur during fall and early winter.

Stable Atmosphere

Unstable Atmosphere

The continent-wide network of weather stations that make regular upper-air soundings gives a broad general picture of the atmospheric structure over North America. These soundings show the major pressure, temperature, and moisture patterns that promote stability, instability, or subsidence, but they frequently do not provide an accurate description of the air over localities at appreciable distances from the upper-air stations. We need, therefore, to supplement these observations with local measurements or with helpful indicators. Clouds, windflow characteristics, occurrence of dust devils, and other phenomena can be useful indicators of stability.

At times, it may be possible to take upper-air observations with portable instruments in fixed-wing aircraft or helicopters. In mountainous country, temperature and humidity measurements taken at mountaintop and valley-bottom stations provide reasonable estimates of the lapse rate and moisture conditions in the air layer between the two levels. In areas where inversions form at night, similar measurements indicate the strength of the inversion. The heights of surface or low-level inversions can be determined by traversing slopes that extend through them. The height at which rising smoke flattens out may indicate the base of a low-level inversion. The tops of clouds in the marine layer along the Pacific coast coincide with the base of the subsidence inversion. The height of the cloud tops provides a good estimate of the height of the inversion.

Other visual indicators are often quite revealing. Stability in the lower layers is indicated by the steadiness of the surface wind. A steady wind is indicative of stable air. Gusty wind, except where mechanical turbulence is the obvious cause, is typical of unstable air. Dust devils are always indicators of instability near the surface. Haze and smoke tend to hang near the ground in stable air and to disperse upward in unstable air.

Cloud types also indicate atmospheric stability at their level. Cumulus-type clouds contain vertical currents and therefore indicate instability. The heights of cumulus clouds indicate the depth and intensity of the instability. The absence of cumulus clouds, however, does not necessarily mean that the air is stable. Intense summer heating can produce strong convective currents in the lower atmosphere, even if the air is too dry for condensation and cloud formation. Generally, though, the absence of clouds is a good indication that subsidence is occurring aloft. Even if scattered cumulus clouds are present during the day and are not developing vertically to any great extent, subsidence very likely is occurring above the cumulus level. Stratus-type cloud sheets indicate stable layers in the atmosphere.

In mountainous country, where fire lookouts on high peaks take observations, a low dew-point temperature may provide the only advance warning of subsidence. Hygrothermograph records and wet- and dry-bulb temperature observations show a sharp drop in relative humidity with the arrival of subsiding air at the mountaintop. Early morning dew-point temperatures of 20° F or lower in summer or early fall may signal the presence of subsiding air, and provide a warning of very low humidities at lower elevations in the afternoon.

Air that rises in the troposphere must be replaced by air that sinks and flows in beneath that which rises. Local heating often results in small-scale updrafts and downdrafts in the same vicinity. On a larger scale, such as the upflow in low-pressure systems, adjacent face high-pressure systems with their divergent flow normally supply the replacement air. The outflow at the surface from these high-pressure areas results in sinking of the atmosphere above them. This sinking from aloft is the common form of subsidence.

The sinking motion originates high in the troposphere when the high-pressure systems are deep. Sometimes these systems extend all the way from the surface up to the tropopause. Deep high-pressure systems are referred to as warm Highs, and subsidence through a deep layer is characteristic of warm Highs.

Subsidence occurs in these warm high-pressure systems as part of the return circulation compensating for the large upward transport of air in adjacent low-pressure areas. If the subsidence takes place without much horizontal mixing, air from the upper troposphere may reach the surface quite warm and extremely dry. For example, the saturation absolute humidity of air in the upper troposphere with a temperature of -50° to -60° F is less than 0.02 pounds per 1,000 cubic feet. In lowering to the surface, this air may reach a temperature of 70° F or higher, where saturation would represent 1.15 pounds or more of water per 1,000 cubic feet. If no moisture were added to the air in its descent, the relative humidity would then be less than 2 percent.

Subsiding air may reach the surface at times with only very little external modification or addition of moisture. Even with considerable gain in moisture, the final relative humidity can be quite low. The warming and drying of air sinking adiabatically is so pronounced that saturated air, sinking from even the middle troposphere to near sea level, will produce relative humidities of less than 5 percent. Because of the warming and drying, subsiding air is characteristically very clear and cloudless.

Subsidence in a warm high-pressure system progresses downward from its origin in the upper troposphere. In order for the sinking motion to take place, the air beneath must flow outward, or diverge. Thus, horizontal divergence is an integral part of subsidence in the troposphere. The descent rate is observed by following the progress of the subsidence inversion on successive upper-air soundings.

The rate of descent of subsiding air varies widely. It is typically fastest at higher levels and becomes progressively slower near the surface. It is commonly about 5,000 feet in 6 hours around the 30,000-foot level, and about 500 feet in 6 hours at the 6,000-foot level.

Frequently, the subsiding air seems to lower in successive stages. When this happens, a sounding will show two or more inversions with very dry air from the top down to the lowest inversion. This air may be drier than can be measured with standard sounding equipment.

Examples of Common Subsidence Patterns

Subsiding air seldom reaches the surface as a broad layer. Often, it sinks to the lower troposphere and then stops. We need, therefore, to consider ways in which the dry air no longer lowering steadily over a broad area can affect the surface.

Subsidence Inversions along the West Coast

Along the west coast in summer we generally find a cool, humid advected marine layer 1,000-2,000 feet thick with a warm, dry subsiding layer of air above it. This subsidence inversion is usually low enough so that coastal mountains extend up into the dry air. The higher topographic elevations will experience warm temperatures and very low humidities both day and night. Some mixing of moisture upward along the slopes usually occurs during the daytime with upslope winds.

As the marine layer moves inland from the coast during clear summer days, it is subjected to intensive heating and becomes warmer and warmer until finally the subsidence inversion is wiped out. The temperature lapse rate from the surface to the base of the dry air, or even higher, becomes dry-adiabatic. Then, convective currents can be effective in bringing dry air from aloft down to the surface and mixing the more moist air from near the surface to higher levels.

This process can well take place in other regions when the subsidence inversion reaches low-enough levels so it can be eliminated by surface daytime heating. The inversion will be wiped out only in local areas where surface heating is intense enough to do the job. If the heating is not sufficient to eliminate the inversion, the warm, dry air cannot reach the surface by convection. Convective currents in the layer beneath the inversion may be effective in eating away the base of the inversion and mixing some of the dry air above with the more humid air below. This process will warm and dry the surface layer somewhat, but humidities cannot reach the extremely low values characteristic of a true subsidence situation.

Sloping

Another method by which dry, subsiding air may reach the surface is by following a sloping downward path rather than a strictly vertical path. A vertical sounding may show that the subsiding air is much too warm to reach the surface by sinking vertically, because the layer beneath it is cooler and denser. However, if surface air temperatures are warmer downstream, the subsiding air can sink dry-adiabatically to lower levels as it moves downstream and may eventually reach the surface. This process is most likely to occur around the eastern and southern sides of a high-pressure area where temperatures increase along the air trajectory. By the time the sinking air reaches the surface, it is likely to be on the south, southwest, or even west side of the High.

Mountain Waves and Foehn Winds

Subsiding air may reach the surface in a dynamic process through the formation of mountain waves when strong winds blow at right angles to mountain ranges. Waves of quite large amplitude can be established over and on the leeward side of ranges. Mountain waves can bring air from great heights down to the surface on the lee side with very little external modification. These waves may also be a part of the foehn-wind patterns.

In the mountain areas of the West, foehn winds, whether they are the chinook of the eastern slopes of the Rockies, the Santa Ana of southern California, or the Mono and northeast wind of central and northern California, are all associated with a high-pressure area in the Great Basin. A foehn is a wind flowing down the leeward side of mountain ranges where air is forced across the ranges by the prevailing pressure gradient.

Subsidence occurs above the High where the air is warm and dry. The mountain ranges act as barriers to the flow of the lower layer so that the air crossing the ranges comes from the drier layer aloft. If the pressure gradient is favorable for removing the surface air on the leeward side of the mountain, the dry air from aloft is allowed to flow down the lee slopes to low elevations. The dryness and warmth of this air combined with the strong wind flow produce the most critical fire-weather situations known anywhere.

Mountain waves, most common and strongest in the West, are also characteristic of flow over eastern and other mountain ranges. When they occur with foehn winds, they create a very spotty pattern. The strongest winds and driest air are found where the mountain waves dip down to the surface on the leeward side of the mountains.

Regional High Pressure Systems

Cases of severe subsidence are much more frequent in the western half of the country than in the eastern regions. Most of the Pacific coast area is affected in summer by the deep semipermanent Pacific High. This provides a huge reservoir of dry, subsiding air which penetrates the continent in recurring surges to produce long periods of clear skies and dry weather. Fortunately, marine air persists much of the time in the lower layer along the immediate coast and partially modifies the subsiding air before it reaches the surface.

In the fall and winter months, the Great Basin High is a frequent source of subsiding air associated with the foehn winds, discussed above. It is the level of origin of this air that gives these winds their characteristic dryness.

Subsiding air reaching the surface is perhaps less common in eastern regions, but does occur from time to time. Usually the subsiding air is well modified by convection. But subsidence is often a factor in the severe fire weather found around the periphery of Highs moving into the region east of the Rockies from the Hudson Bay area or Northwest Canada mostly in spring and fall. It also occurs during summer and early fall periods of drought, when the Bermuda High extends well westward into the country.

General winds are produced by the broadscale pressure gradients which are shown on synoptic weather maps, but may be modified considerably by friction or other topographic effects. The amount of influence of the surface on general winds is largely dependent on wind speed and the stability of the air. Stable air flowing over even surfaces tends to be smooth, or laminar. Unstable air or strong winds flowing over rough surfaces is turbulent and full of eddies. Forests and other vegetated areas are characteristically rough surfaces and thus contribute to air turbulence, eddies, etc.

Surface winds in the Northern Hemisphere tend to shift clockwise with the passage of fronts. In mountainous topography, however, the effect of the mountains on the windflow usually overshadows this. The windflow is channeled, and, over sharp crests, eddies are produced. At times, waves form over mountains, and, if conditions are favorable, strong surface winds are experienced on the lee side. When the airflow is from higher to lower elevations, the air warms adiabatically and foehn winds are produced.

Wind affects wildfire in many ways. It carries away moisture-laden air and hastens the drying of forest fuels. Light winds aid certain firebrands in igniting a fire. Once a fire is started, wind aids combustion by increasing the oxygen supply. It aids fire spread by carrying heat and burning embers to new fuels, and by bending the flames closer to the unburned fuels ahead of the fire. The direction of fire spread is determined mostly by the wind direction. Thus the fire-control plan, in the case of wildfire, and the burning plan, in the case of prescribed fire, must be based largely on the expected winds.

• Laminar Flow
• Mechanical and Thermal Turbulence
• Effects of Vegetation on Surface Winds
• Frontal Winds
• Effects of Mountain Topography on Surface Winds
• Mountain Waves
• Foehn Winds
• Winds Aloft
• Measuring Surface Winds

The absence of turbulence in surface winds--a steady even flow--is called laminar flow. The term suggests air moving along in flat sheets or layers, each successive thin layer sliding over the next. Laminar or near-laminar flow occurs in stable air moving at low speeds. It is characteristic of cold air flowing down an incline, such as we might find in a nighttime inversion. The air flows smoothly along, following the topography and varying little in speed. Vertical mixing is negligible.

True laminar flow is probably rare in wildland fire situations, but, on occasion, turbulence is minor and, for all practical purposes, surface winds do have the steady speed and direction characteristic of laminar motion. While turbulent winds usually cause more erratic fire behavior, the laminar type may result in more rapid and sustained fire spread in one direction. Laminar flow is most likely to occur at night. It is frequently observed over open plains and gently rolling topography.

Surface winds often vary considerably in both speed and direction over short intervals of time. They tend to blow in a series of gusts and lulls with the direction fluctuating rapidly. This irregular air motion is known as turbulence, which may be either mechanical or thermal in nature. At the surface, turbulence is commonly identified in terms of eddies, whirls, and gusts; aloft it is associated with "bumpy" flying.

The depth of the air layer through which the frictional force is effective also varies with the roughness of the surface; it is shallower over smooth surfaces and deeper over rough topography. The depth may also vary with the stability of the lower atmosphere. A low inversion will confine the frictional effect to a shallow surface layer, but a deep layer can be affected if the air is relatively unstable. These effects vary widely both with time and between localities. Usually the friction layer is considered to be about 2,000 feet deep. The top of the friction layer is the gradient wind level above which the windflow tends to parallel the isobars or pressure-surface contours.

Mechanical turbulence

Turbulent Flow

Surface friction produces mechanical turbulence in the airflow. The effect of friction is least over smooth water and greatest over mountainous topography. The flow of stable air near the surface is similar to the flow of water in a creekbed. At low speeds the currents of air tend to follow the general contours of the landscape. But when the speed increases--as when a creek rises--the current "tumbles" over and around hills and ridges, structures, trees, and other obstacles, and sets up eddies in all directions (for example, see Effects of Vegetation on Surface Winds). Mechanical turbulence increases with both wind speed and the roughness of the surface.

Thermal turbulence

Thermal turbulence is associated with instability and convective activity. It is similar to mechanical turbulence in its effects on surface winds, but extends higher in the atmosphere. Since it is the result of surface heating, thermal turbulence increases with the intensity of surface heating and the degree of instability indicated by the temperature lapse rate. It therefore shows diurnal changes, and is most pronounced in the early afternoon when surface heating is at a maximum and the air is unstable in the lower layers. It is at a minimum during the night and early morning when the air is more stable. Mechanical and thermal turbulence frequently occur together, each magnifying the effects of the other.

Thermal turbulence induced by the combination of convection and horizontal wind is the principal mechanism by which energy is exchanged between the surface and the winds aloft. Unstable air warmed at the surface rises to mix and flow along with the winds above. This turbulent flow also brings air with higher wind speeds--greater momentum--from aloft down to the surface, usually in spurts and gusts. This momentum exchange increases the average wind speed near the surface and decreases it aloft. It is the reason why surface winds at most places are stronger in the afternoon than at night.

Eddies

Eddies

Eddy formation is a common characteristic of both mechanical and thermal turbulent flow. Every solid object in the wind path creates eddies on its lee side. The sizes, shapes, and motions of these eddies are determined by the size and shape of the obstacle, the speed and direction of the wind, and the stability of the lower atmosphere. Although eddies may form in the atmosphere with their axes of rotation in virtually any plane, it is usual to distinguish between those which have predominantly vertical or horizontal axes. A whirlwind or dust devil is a vertical eddy, as are eddies produced around the corners of buildings or at the mouths of canyons with steep sides. Large, roughly cylindrical eddies that roll along the surface like tumbleweeds are horizontal eddies.

Eddies associated with individual fixed obstructions tend to remain in a more-or-less stationary position in the lee of the obstruction. If they break off and move downstream, new ones form near the obstruction. The distance downwind that an obstacle, such as a windbreak, affects the windstream is variable. For most obstructions, the general rule of thumb is that this distance is 8 to 10 times the height of the obstacle.

Rotation speeds in eddies are often much greater than the average wind speeds measured with mechanical anemometers. These higher speeds are often of short duration at any point, except where stationary eddies are found, but are still significant in fire behavior. Whirlwinds, for example, develop speeds capable of lifting sizable objects. Eddies moving with the general windflow account for the principal short-term changes in wind speed and direction known as gustiness.

Vegetation is part of the friction surface which determines how the wind blows near the ground. Forests and other vegetated areas are characteristically rough surfaces and thus contribute to air turbulence, eddies, etc. They also have the distinction of being somewhat pervious, allowing some air movement through, as well as over and around, the vegetation.

Wind speeds over open, level ground, although zero at the very surface, increase quite rapidly in the first 20 feet above the ground. Where the surface is covered with low-growing, dense vegetation such as grass or brush, it is satisfactory, for most weather purposes, to consider the effective friction surface as the average height of the vegetation, disregarding the air flowing through it. In areas forested with trees, however, airflow within and below the tree canopies is important.

The leaf canopy in a forest is very effective in slowing down wind movements because of its large friction area. In forests of shade-tolerant species where the canopy extends to near ground level, or in stands with understory vegetation, wind speed is nearly constant from just above the surface to near the tops of the crowns. Above the crowns, wind speed increases much as it does over level ground. In forest stands that are open beneath the main tree canopy, air speed increases with height above the surface to the middle of the trunk space, and then decreases again in the canopy zone.

How much the wind speed is reduced inside the forest depends on the detailed structure of the forest stand and on wind speed above the forest canopy, or as measured out in the open away from the forest. The drag of any friction surface is relatively much greater at high wind speeds than it is with low speeds. At low wind speeds, the forest may have only a small effect on the speed of the wind. For example, a 4-m.p.h. wind measured in the open might be slowed to 2.5-m.p.h. at the same height inside the forest. But a fairly high wind speed in the open will be slowed in the forest in much greater proportion. Thus, a 20-m.p.h. wind might be reduced to 4- or 5-m.p.h. in an 80 foot-tall stand of second-growth pine with normal stocking. The reduction would vary considerably, however, among different species and types of forest. Deciduous forests have a further seasonal variation, because although trees bare of leaves have a significant effect in limiting surface wind speeds, it is far less than when the trees are in full leaf.

Local eddies are common in forest stands and are found in the lee of each tree stem. These small eddies affect the behavior of surface fires. Larger scale eddies often form in forest openings. The higher winds aloft cause the slower moving air in these openings to rotate about a vertical axis, or roll over in a horizontal manner. The surface wind direction is then frequently opposite to the direction above the treetops.

The edges of tree stands often cause roll eddies to form in the same manner as those associated with bluffs. Wind blowing against the stand often produces small transient eddies on the windward side, while those in the lee of a forest are mostly larger and more fixed in location, with subeddies breaking off and moving downwind.

Strong surface heating, as on warm, sunny days, adds to the complexity of these forest airflow patterns. Thermal turbulence is added to the generally turbulent flow through open timber stands as it is to the flow above a closed forest canopy. The flow beneath a dense canopy is affected only slightly by thermal turbulence, except where holes let the sun strike bare ground or litter on the forest floor. These become hotspots over which there is a general upwelling of warm air through the canopy. This rising air is replaced by gentle inflow from surrounding shaded areas. Thermal turbulence on the lee side of a forest stand may often be enough to disguise or break up any roll eddies that tend to form (see Effects of Vegetation on Surface Temperature).

As successive air masses move across the land, the change from one to another at any given point is marked by the passage of a front. A front is the boundary between two air masses of differing temperature and moisture characteristics. The type of front depends upon the movement of the air masses.

Where a cold air mass is replacing a warm air mass, the boundary is called a cold front. Where a warm air mass is replacing a cold air mass, the boundary is called a warm front. If a cold front overtakes a warm front, the intervening warm air is lifted from the surface, and the air mass behind the cold front meets the air mass ahead of the warm front. The frontal boundary between these two air masses is then called an occlusion or occluded front.

Fronts are most commonly thought of in association with precipitation and thunderstorms. But occasionally fronts will cause neither. In these instances, the winds accompanying the frontal passage may be particularly significant to fire behavior. The variability of general surface winds during the spring and fall fire seasons is somewhat greater in eastern portions of the continent than during the summer fire season of the mountainous West. The East experiences more frequent and rapid movement of pressure systems than occur in the West. In the West, the major mountain chains tend both to hinder the movement of Highs and Lows and to lift winds associated with them above much of the topography. Strong summer surface heating also diminishes the surface effects of these changes.

Shifts in wind direction associated with fronts

The passage of a front is usually accompanied by a shift in wind direction. The reason for this is that fronts lie in troughs of low pressure. The isobars in a trough are curved cyclonically in the Northern Hemisphere. This means that as a trough, with its front, passes a particular location the wind direction shifts clockwise. The wind behavior during the frontal passage depends upon the type of front, its speed, the contrast in temperature of the air masses involved, and upon local conditions of surface heating and topography.

Warm fronts

East of the Rockies, the surface wind ahead of a warm front usually blows from a southeasterly or southerly direction. With the frontal passage, the wind gradually shifts clockwise. The change in wind direction usually amounts to between 45° and 90°; therefore, after the warm front goes by, the wind commonly blows from the southwest. Steady winds, rather than gusty winds, both before and after the frontal passage are the rule, because the layer of air next to the ground is generally stable. Warm-front passages in the mountainous West are fewer, more erratic, and tend to become diffuse.

Cold fronts

The passage of a cold front differs from that of a warm front. The wind change is usually sharp and distinct, even when the air is so dry that few if any clouds accompany the front. Ahead of a cold front, the surface wind is usually from the south or southwest. As the front approaches, the wind typically increases in speed and often becomes quite gusty. If cold air aloft overruns warm air ahead of the front at the surface, the resulting instability may cause violent turbulence in the frontal zone. The wind shift with the passage of a cold front is abrupt and may be less than 45° or as much as 180°.

After the front has passed, the wind direction is usually west, northwest, or north. Gustiness may continue for some time after the frontal passage, because the cooler air flowing over warmer ground tends to be unstable. This is particularly true in the spring months. If the temperature contrast is not great, however, the winds soon become steady and relatively gentle.

Occluded fronts

The wind shift accompanying the passage of an occluded front is usually 90° or more. The wind generally shifts from a southerly direction to a westerly or northwesterly direction as the occlusion passes. The wind shift with an occlusion resembles that of a warm front or cold front, depending upon whether the air behind the occlusion is warmer or colder than the air ahead. The violent turbulence that may accompany a cold-front passage, however, is usually absent with an occluded frontal passage.

In the area east of the Rockies, squall lines often precede cold fronts. These are narrow zones of instability that usually form ahead of and parallel to the cold front. Most common in the spring and summer, squall lines are associated with severe lightning storms in the Midwest and may have extremely violent surface winds. They usually develop quickly in the late afternoon or night, move rapidly, and tend to die out during late night or early morning.

Winds ahead of the squall are usually from a southerly direction. They increase to 30, 40, or even 60 miles per hour, shift to the west or northwest, and become extremely gusty as the squall line passes. The strong, gusty winds ordinarily do not last long, and the winds soon revert to the speed and direction they had prior to the squall. This wind behavior distinguishes a squall line from a cold front. Squall lines are usually accompanied by thunderstorms and heavy rain. But occasionally the storms are scattered along the line so that any one local area might experience squall-line wind behavior without the fire-quenching benefit of heavy rain.

Mountains represent the maximum degree of surface roughness and thus provide the greatest friction to the general surface airflow. Mountain chains are also effective as solid barriers against airflow--particularly dry, cold air of polar origin and relatively cool Pacific marine air. While warm, light air may be forced aloft and flow over the ranges, cool, heavy air is often either dammed or deflected by major mountain systems.

Over short distances and rough topography, gradient balance may not be established and winds of considerable speed may blow almost directly across isobars from higher to lower pressure. Winds of this nature are common in both coastal and inland mountain regions. This type of flow is particularly noticeable in the strong pressure-gradient region of a Santa Ana pattern.

Mountains and their associated valleys provide important channels that establish local wind direction. Airflow is guided by the topography into the principal drainage channels. Less-prominent features of the landscape have similar, though smaller scale, local mechanical effects on wind speed, direction, and turbulence. In short, winds blowing over the surface are influenced by every irregularity.

In addition to these mechanical effects, strong daytime convective activity in mountain areas often alters or replaces the general wind at the surface. General winds are most pronounced at the surface in the absence of strong heating.

Canyon Eddies

General winds blowing across mountain ridges are lifted along the surface to the gaps and crests. If the air is stable, it will increase in speed as it crosses the ridge. Ridgetop winds thus tend to be somewhat stronger than winds in the free air at the same level. How the air behaves on crossing a ridge is influenced by ridge shape and wind speed and direction. Round-topped ridges tend to disturb surface airflow the least. In light to moderate winds there is often little evidence of any marked turbulence. Sharp ridges, on the other hand, nearly always produce significant turbulence and numerous eddies on the lee side. Some of this is evident at the surface as gusts and eddies for short distances below the ridgetop, though much of it continues downwind aloft. Wind blowing perpendicular to the ridge line develops the least complex wind structure downwind, and most of the eddies formed are of the roll or horizontal type. If the angle of wind approach deviates from the perpendicular by some critical amount, perhaps 30° or less, vertical eddies are likely to be found in the lee draws below the ridgetop, in addition to eddies in other planes.

Eddy currents are often associated with bluffs and similarly shaped canyon rims. When a bluff faces downwind, air on the lee side is protected from the direct force of the wind flowing over the rim. If the wind is persistent, however, it may start to rotate the air below and form a large, stationary roll eddy. This often results in a moderate to strong upslope wind opposite in direction to that flowing over the rim. Eddies of this nature are common in the lee of ridges that break off abruptly, and beneath the rims of plateaus and canyon walls.

Ridgetop saddles and mountain passes form important channels for local pressure-gradient winds. Flow converges here as it does across ridgetops, with an accompanying increase in wind speed. After passing through mountain saddles, the wind often exhibits two types of eddy motion on the lee side. One takes the form of horizontal eddies rolling or tumbling down the lee slope or canyon, although the main eddy may be stationary. The other is usually a stationary vertical eddy in one of the sheltered areas on either side of the saddle. Some of these vertical eddies may also move on downwind.

General winds that are channeled in mountain canyons are usually turbulent. The moving air in canyons is in contact with a maximum area of land surfaces. Alternating tributaries and lateral ridges produce maximum roughness. Whether the canyon bottom is straight or crooked also has an important influence on the turbulence to be expected. Sharp bends in mountain-stream courses are favorite "breeding grounds" for eddies, particularly where the canyon widens to admit a side tributary. Such eddies are most pronounced near the canyon floor and dissipate well below the ridgetop.

See: Mountain Waves and Foehn Winds

Mountain Waves

Moderate to strong winds in a stably stratified atmosphere blowing across high mountain ranges will cause large-scale mountain waves for many miles downwind. The stable air, lifted by the wind over the mountain range, is pulled downward by gravity on the lee side. Inertia carries the air past its equilibrium level, so it rises again farther downslope. This oscillatory motion forms a series of lesser waves downstream until the oscillation finally ceases. Waves may extend as high as 40,000 feet or more in the well-known Bishop wave in California. Large-scale waves occur in the Rocky Mountains, and waves on a lesser scale appear in the Appalachians and elsewhere.

The lee slope of the mountains may experience strong downslope winds or many eddies of various sizes which roll down the slope. Within each wave downstream from the mountain range, a large roll eddy may be found with its axis parallel to the mountain range. Roll eddies tend to be smaller in each succeeding wave downstream. The waves downwind of the mountains are referred to as lee waves or standing waves.

If sufficient moisture is present, cap clouds will form over the crest of the mountains, roll clouds will be found in the tops of the roll eddies downstream, and wave clouds will be located in the tops of the waves.

Foehn winds represent a special type of local wind associated with mountain systems, particularly in the Western United States. In most mountainous areas, local winds are observed that blow over the mountain ranges and descend the slopes on the leeward side. If the downflowing wind is warm and dry, it is called a foehn wind. The wind is called a bora or fall wind if the air is originally so cold that even after it is warmed adiabatically in flowing down the mountain slopes it is still colder than the air it is replacing on the leeward side. The bora rarely occurs in North America and is not important in this discussion, because of its cold temperatures and the fact that the ground is often snow-covered when it occurs. We are concerned more with the warmer foehn, which creates a most critical fire-weather situation.

The development of a foehn wind requires a strong high-pressure system on one side of a mountain range and a corresponding Low or trough on the other side. Such pressure patterns are most common to the cool months; therefore, foehn winds are more frequent in the period from September through April than in the summer months.

Two types of foehn winds are common in mountains of the Western U.S.

Foehn winds of the first type result when a deep layer of moist air is forced upward and across a mountain range. As the air ascends the windward side, it is cooled dry-adiabatically until the condensation level is reached. Further lifting produces clouds and precipitation, and cooling at the lesser moist-adiabatic rate. The water vapor that has condensed and fallen out as precipitation is lost to the air mass. Upon descending the leeward slopes, the air mass warms first at the moist-adiabatic rate until its clouds are evaporated. Then it warms at the dry-adiabatic rate and arrives at lower elevations both warmer and drier than it was at corresponding levels on the windward side. In descending to the lowlands on the leeward side of the range, the air arrives as a strong, gusty, desiccating wind.

Moist Pacific air forced across the Sierra Cascade range loses some of its moisture and exhibits mild foehn characteristics on the eastern slopes. Forced across the Rocky Mountain range, the same air loses additional moisture and may produce a well-developed foehn on the eastern slopes in that region. The Plains east of the Rockies are often under the influence of a cold air mass of Canadian origin in the cooler months. If this air mass is then moved eastward by a favorable pressure gradient and replaced by a warm descending foehn, abrupt local temperature rises are experienced.

The second type of foehn is related to a cold, dry, usually stagnated high-pressure air mass restricted by mountain barriers. If a low-pressure center or trough is located on the opposite side of the barrier, the strong pressure gradient will cause air to flow across the mountains. Since the mountains block the flow of surface air, the airflow must come from aloft. The air above the surface high-pressure system is subsiding air and is therefore dry and potentially quite warm. On the leeward side of the mountains, surface air is forced away by the strong pressure gradient, and it is replaced by the air flowing from aloft on the windward side and descending to the lowland on the leeward side. Surface wind speeds of 40 to 60 miles per hour are common in foehn flow of this type, and speeds up to 90 miles per hour have been reported. The wind often lasts for 3 days or more, with gradual weakening after the first day or two. Sometimes, it stops very abruptly.

High-pressure areas composed of cool air masses frequently stagnate in the Great Basin of the Western United States during the fall, winter, and spring months. Depending on its location, and the location of related Lows or troughs, a Great Basin High may create foehn winds which move eastward across the northern and central Rockies, westward across the Oregon and Washington Cascades and the northern and central Sierra Nevada, or southwestward across the Coast Ranges in southern California. A combination of high pressure over the State of Washington and low pressure in the Sacramento Valley causes north winds in northern California. Brief foehn wind periods, lasting 1 or 2 days, may result from migrating Highs passing through the Great Basin.

The course of the foehn may be either on a front many miles wide or a relatively narrow, sharply defined belt cutting through the leeside air, depending on the pressure pattern and on the topography.

A foehn, even though it may be warm, often replaces cooler air on the lee side of the mountains. Counterforces sometimes prevent this, however, and cause the foehn to override the cooler air and thus not be felt at the surface at lower elevations. At other times the foehn may reach the surface only intermittently, or at scattered points, causing short-period fluctuations in local weather.

Two mechanisms come into play. One is a favorable pressure gradient acting on the lee-side air in such a way as to move it away from the mountains so that the warm foehn can replace it.

A second mechanism is the mountain wave phenomenon. The wavelength and wave amplitude depend upon the strength of the flow bearing against the mountains and the stability of the layers in which the wave may be embedded . When these factors are favorable for producing waves which correspond to the shape of the mountain range, the foehn flow will follow the surface and produce strong surface winds on the lee slopes. There is evidence that strong downslope winds of the warm foehn on lee slopes are always caused by mountain waves. The change in wavelength and amplitude can account for the observed periodic surfacing and lifting of foehn flow. Surfacing often develops shortly after dark as cooling stabilizes the air crossing the ridge.

Foehn winds in the Western United states

The Chinook

The Chinook, a foehn wind on the eastern slopes of the Rocky Mountains, often replaces cold continental air in Alberta and the Great Plains. Quick wintertime thawing and rapid snow evaporation are characteristic. If the cold air is held in place by the local pressure and circulation system, the foehn will override it; or if the cold air stays in the bottoms because of its greater density, the Chinook may reach the surface only in the higher spots. Relative humidities dropping to 5 percent or less and temperature changes of 30° F. to 40° F. within a few minutes are common in Chinooks.

East winds

Along the Pacific coast a weak foehn may be kept aloft by cool marine air flowing onshore. On the other hand, a strong, well-developed foehn may cut through all local influences and affect all slope and valley surfaces from the highest crest to the sea. East winds in the Pacific Northwest, for example, sometimes flow only part way down the lee slopes of the Cascades, and then level off above the lowlands and strike only the higher peaks and ridges of the coastal mountains. At other times virtually all areas are affected.

North and Mono winds

North and Mono winds in northern and central California develop as a High moves into the Great Basin. North winds develop if a High passes through Washington and Oregon while a trough is located in the Sacramento Valley. Mono winds occur after the High has reached the Great Basin, providing there is a trough near the coast. Both North and Mono are foehn winds bringing warm, dry air to lower elevations. At times they will affect only the western slopes of the Sierra Nevada, and at other times they push across the coastal mountains and proceed out to sea. This depends upon the location of the low-pressure trough. These winds are most common in late summer and fall.

Santa Ana winds

The Santa Anas of southern California also develop with a High in the Great Basin. The low-pressure trough is located along the southern California coast, and a strong pressure gradient is found across the southern California mountains.

In the coastal mountains, and the valleys, slopes, and basins on the ocean side, the Santa Ana varies widely. It is strongly channeled by the major passes, and, at times, bands of clear air can be seen cutting through a region of limited visibility. The flow coming over the tops of the ranges may remain aloft on the lee side or drop down to the surface, depending upon whether the Santa Ana is "strong" or "weak" and upon its mountain-wave characteristics. If the foehn flow is weak and remains aloft, only the higher elevations in the mountains are affected by the strong, dry winds. Local circulations, such as the sea breeze and slope winds, are predominant at lower elevations, particularly in areas away from the major passes.

Typically in southern California during the Santa Ana season, there is a daytime onshore breeze along the coast and gentle to weak upslope and upcanyon winds in the adjacent mountain areas. With nighttime cooling, these winds reverse in direction to produce downcanyon and offshore winds, usually of lesser magnitude than the daytime breeze. A strong Santa Ana wind wipes out these patterns. It flows over the ridges and down along the surface of leeward slopes and valleys and on to the sea. The strong winds, along with warm temperatures and humidities sometimes lower than 5 percent, produce very serious fire weather in a region of flashy fuels. The strong flow crossing the mountains creates mechanical turbulence, and many eddies of various sizes are produced by topographic features.

A strong Santa Ana, sweeping out the air ahead of it, often shows little or no difference in day and night behavior in its initial stages. But, after its initial surge, the Santa Ana begins to show a diurnal behavior. During the daytime, a light sea breeze may be observed along the coast and light upvalley winds in the coastal valleys. The Santa Ana flow is held aloft, and the mountain waves are not of proper dimensions to reach the surface. The air in the sea breeze may be returning Santa Ana air, which has had only a short trajectory over the water and is not as moist as marine air. After sunset, the surface winds reverse and become offshore and downslope. Increasing air stability may allow the shape of the mountain waves to change so that the lower portions of waves can strike the surface and produce very strong winds down the lee slopes. As the Santa Ana continues to weaken, the local circulations become relatively stronger and finally the normal daily cycle is resumed.

Wind Profiles 1 & 2

Wind Profiles 3 & 4

Usually, we separate winds into surface winds and winds aloft. There is no sharp separation between them, but rather a blending of one into the other. We think of surface winds as those winds measured with instruments mounted on surface-borne masts or towers. Winds aloft are those measured with airborne equipment from the surface layer up to the limit of our interest. In ascending from the surface through the lower atmosphere, there is a transition in both speed and direction from the surface to the top of the friction layer, which is also called the mixing layer. The depth of this friction or mixing layer is dependent upon the roughness of the terrain and the intensity of heating or cooling at the surface. The winds aloft above the mixing layer are steader in speed and direction, but they do change as pressure centers move and change in intensity.

Pressure systems higher in the troposphere may differ markedly from those near the surface. At progressively higher altitudes, closed pressure systems are fewer. Furthermore, it is common for the troposphere to be stratified or layered. With height, there may be gradual changes in the distribution of Highs and Lows. These changes produce different wind speeds and directions in the separate layers. With strong stratification the wind direction may change abruptly from one layer to the next. The difference in direction may be anywhere from a few degrees to complete reversal. In the absence of marked stratification above the friction layer, wind direction at adjacent levels tends to be uniform, even though the speed may change with altitude. A common cause of stratification in the lower troposphere is the overriding or underrunning of one air mass by another. Thus, the layers often differ in temperature, moisture, or motion, or in any combination of these.

Marked changes in either wind speed or direction between atmospheric layers often occur with an inversion which damps or prevents vertical motion, whether it is convection over a fire or natural circulation in the formation of cumulus clouds. Even though a wind speed profile--a plot of wind speed against height--of the upper air might indicate only nominal air speeds, the relative speeds of two air currents flowing in nearly opposite directions may produce strong wind shear effects. Wind shear in this case is the change of speed or direction with height. Clouds at different levels moving in different directions, tops being blown off growing cumulus clouds, and rising smoke columns that break off sharply and change direction are common indicators of wind shear and disrupted vertical circulation patterns.

Wildland fires of low intensity may be affected only by the airflow near the surface. But when the rate of combustion increases, the upper airflow becomes important as an influence on fire behavior. Airflow aloft may help or hinder the development of deep convection columns. It may carry burning embers which ignite spot fires some distance from the main fire. The winds aloft may be greatly different in speed and direction from the surface winds.

Wind is air in motion relative to the earths surface. Its principal characteristics are its direction, speed, and gustiness or turbulence. Wind direction and speed are usually measured and expressed quantitatively, while in field practice turbulence is ordinarily expressed in qualitative or relative terms. Ordinarily only the horizontal components of direction and speed are measured and reported, and this is adequate for most purposes. In fire weather, however, we should remember that winds can also have an appreciable vertical component which will influence fire behavior, particularly in mountainous topography.

At weather stations making regular weather observations, surface wind direction is determined by a wind vane mounted on a mast and pointing into the wind. The direction can be determined visually or, with more elaborate instruments, it can be indicated on a dial or recorded on a chart.

Wind direction is ordinarily expressed as the direction from which the wind blows. Thus, a north wind blows from the north toward the south, a northeast wind from the northeast, and so on around the points of the compass. Direction is also described in degrees of azimuth from north--a northeast wind is 45°, a south wind 180°, and a northwest wind 315°.

The method of describing the direction of both surface winds and winds aloft, by the direction from which the wind blows, is ordinarily very practical. In mountain country, though, surface wind direction with respect to the topography is often more important in fire control and provides a better description of local winds than the compass direction. Here it is common to express the wind direction as the direction toward which the wind is headed. Thus, an upslope or upcanyon wind is actually headed up the slope or up the canyon. Wind is described as blowing along the slopes, through the passes, or across the ridges. Similarly, "offshore" or "onshore" are used to describe the directions toward which land and sea breezes are blowing.

Surface wind speeds are measured with anemometers. Many types of anemometers are in use, but the most common is the cup anemometer. It indicates either the air speed at any given instant or the miles of air that pass the instrument in a given time period. The latter gives an average wind for the selected time period. Normally, a 2-minute average is used. The standard height at which wind speed is measured is 20 feet above open ground.

In the United States, wind speed is usually measured in miles per hour or knots (nautical miles per hour). One knot is 1.15 miles per hour. Weather Bureau and military weather agencies use knots for both surface and upper winds, while miles per hour is still in common use in many other agencies and operations, including fire weather.

The direction and speed of winds aloft are determined most commonly by tracking an ascending, gas-filled balloon from the surface up through the atmosphere.

The simplest system employs a pilot balloon followed visually with a theodolite. If a constant rate of rise of the balloon is assumed, periodic readings of elevation and azimuth angles with the theodolite allow computation of average wind direction and speed between balloon positions. Errors are introduced when the ascent rate is not constant because of vertical air currents. If a radiosonde unit (which transmits temperature, moisture, and pressure data during ascent) is added to the balloon, the height of the balloon at the time of each reading can be calculated fairly accurately, and the computed winds are more accurate.

The most refined of present systems have the further addition of a self-tracking, radio direction-finding unit that measures elevation and azimuth angles, and slant range from the observing station to the balloon. This unit known as a rawinsonde, yields quite accurate upper-air information. All of these methods furnish wind soundings for meteorological use and interpretation.

The speed and direction of upper winds are sampled at regular intervals each day at selected weather stations across the continent. These stations are often more than 100 miles apart. Although winds aloft tend to be more uniform than surface winds, there are exceptions. The wind structure over an area some distance from a sampling station may differ considerably from that indicated by the nearest sounding.

Convective winds originate from small-scale pressure gradients produced by local temperature differences. Air made buoyant by warming at the surface is forced aloft; air which is cooled tends to sink. Convective winds may be strengthened, weakened, or eliminated by general winds. The influence of these general winds on the convective wind systems varies with the strength of the general wind, its direction relative to the convective circulation, and the stability of the lower atmosphere.

The nature and strength of convective winds vary with many other factors. Since they are temperature-dependent, all features of the environment that affect heating and cooling are significant. Among the more important are season, diurnal changes, cloud cover, nature of the terrain and its cover such as water, vegetation, or bare ground, and the moisture and temperature structure of the overlying atmosphere.

Convective Winds

The most familiar convective winds are land and sea breezes, valley and slope winds, whirlwinds, and winds associated with convective cumulus and thunderstorm clouds. In the land- and sea-breeze system, the local winds are due to land-water temperature differences, which, in turn, produce differences in the temperature of the overlying air. Slope winds are due to temperature differences between slope air and air over the valley. Valley winds likewise result from temperature differences between valley air and air at the same elevation over the plains. Strong local heating will develop a very unstable layer of air near the surface, and the sudden release of this concentrated energy, usually following a triggering action, may produce whirlwinds. Thermal updrafts resulting from local heating may produce cumulus clouds, which, under suitable moisture and instability conditions, may develop into thunderstorms. Updrafts are convective winds characteristic of developing cumulus clouds, but downdrafts are produced in thunderstorms after precipitation begins falling from the cloud.

• Land and Sea Breezes
  • Gulf and Atlantic Breezes
  • Pacific Coast Sea Breeze
• Slope and Valley Winds
  • Slope Winds
  • Valley Winds
  • Effects of Orientation and Vegetation
  • Interaction of Valley and Slope Winds With General Winds
• Whirlwinds
  • Firewhirls
• Thunderstorm Winds

During the daytime, when land surfaces become warmer than adjacent water surfaces, the air over the land expands, becomes less dense, and the pressure becomes lower than that over the nearby water. As a result of this local-scale pressure difference, a sea breeze begins to flow inland from over the water, forcing the warm air over the land to rise and cool adiabatically. In the absence of strong general winds, this air flows seaward aloft to replace air which has settled and moved toward shore, and thus completes the circulation cell.

Land/Sea Breeze

The surface sea breeze begins around midforenoon, strengthens during the day, and ends around sunset, although the times can vary considerably because of local conditions of cloudiness and the general winds. The breeze begins at the coast, then gradually pushes farther and farther inland during the day, reaching its maximum penetration about the time of maximum temperature.

The land breeze at night is the reverse of the daytime sea-breeze circulation. At night, land surfaces cool more quickly than water surfaces. Air in contact with the land then becomes cooler than air over adjacent water. The increase in air density causes pressure to become relatively higher over the land than over the water, and this pressure difference, in turn, causes air to flow from the land to the water. The air must be replaced, but any return flow aloft is likely to be so weak and diffuse that it is lost in the prevailing general winds.

The land breeze begins 2 to 3 hours after sunset and ends shortly after sunrise. It is a more gentle flow than the sea breeze, usually about 3 to 5 miles per hour. The land air, having been cooled from below by contact with the ground, is stable. The land breeze is, therefore, more laminar and shallower than the sea breeze.

The daily land and sea breezes tend to occur quite regularly when there is no significant influence from the general wind flow. When general winds are sufficiently strong, however, they usually mask the land and sea breezes. A general wind blowing toward the sea opposes the sea breeze and, if strong enough, may prevent its development. In any case the sea breeze is delayed. Depending on the strength of the general wind, this delay may extend into the afternoon. This often results in a "piling up" of marine air off the coast. Then, when the local pressure difference becomes great enough, this sea air moves inland with the characteristics of a small-scale cold front. Air behind the front is initially cool and moist but warms rapidly as it moves over sun-warmed land.

The land breeze does not form against a strong onshore general wind. It is common, however, for the land breeze to slide under onshore winds of light speeds. In doing so, the land breeze does not extend very far seaward.

General winds, either in the direction of the land or sea breeze, or parallel to the coast, tend to mask the true land- or sea-breeze component. Strong general winds produce mechanical mixing which tends to lessen the temperature difference between the land and the sea surfaces. Thus the sea-breeze component becomes weak and only slightly alters the general wind flow. General winds also tend to mask out the closed-cell feature of the land- and sea-breeze circulations by overshadowing the return flow aloft. With an onshore general wind aloft, for example, there is no return flow in the daytime sea-breeze circulation.

General winds along an irregular coastline may oppose a land or sea breeze in one sector and support it in another. Oftentimes, too, shifting general winds may cause periodic reversals of these effects in nearby localities, and may result in highly variable local wind patterns.

Land and sea breezes occur along much of the Pacific coast, the Gulf of Mexico, and the Atlantic seaboard. Gulf and Atlantic breezes and Pacific Coast sea breezes differ in their respective behaviors due to marked differences in general circulation patterns, temperature contrasts, and topography. Whether or not these factors are significant locally depends on the local climate and on the shape and orientation of the shoreline and inland topography.

In the East, land and sea breezes are most pronounced in late spring and early summer, when land and water temperature differences are greatest, and they taper off toward the end of the warm season as temperature differences decrease. They are sufficiently strong during the spring and fall fire seasons to warrant consideration as important fire-weather elements in coastal areas.

Land- and sea-breeze circulations in the East are more often dominated by changes in the general wind pattern than they are in the West. Otherwise, the eastern land and sea breeze represents a more simple situation than the western because coastal topography is flat and uniform.

During the fire season in the East, general circulation patterns are such that on both the Gulf and Atlantic shores there are frequent periods of onshore or offshore winds strong enough to block or mask out land- and sea-breeze development. Onshore general winds almost always mask sea-breeze effects. During periods of gentle to moderate offshore winds, on the other hand, the sea breeze may develop and move inland. Against an opposing general wind, however, the sea breeze moves forward behind a small-scale cold front. This moves slowly, perhaps 3 or 4 miles an hour, and at times may oscillate back and forth with the varying force of the general wind. In addition to the rapid changes in wind speed and direction associated with a cold-front passage, a small area may thus be subjected to several of these passages over a considerable time. At this slow and intermittent pace, the sea breeze may have penetrated inland only a few miles by late afternoon.

Another feature of this type of sea breeze is that it is operating in an area of convergence. This is conducive to turbulent vertical motion in addition to the above-mentioned horizontal surface disturbances. This combination can create critical fire-weather situations, particularly in view of the fact that this type of sea breeze is prone to occur on high fire-danger days.

The reverse land breeze often becomes just part of the offshore general wind and thereby loses its identity.

The Pacific, coastal area sea breeze is at its peak at the height of the summer fire season. It is an important feature of the summer weather along much of the Pacific coast. Water temperatures there are much lower than along the Gulf of Mexico and the Atlantic coast. Intense daytime land heating under clear skies is an additional factor in producing greater land-water temperature differences along the Pacific coast. The sea breeze is, therefore, stronger along the western than the eastern coasts. It is a daily summertime occurrence along the Pacific coast except on rare occasions when it is opposed by the general circulation.

Normally the general wind serves to strengthen the Pacific coast sea breeze. During the summer months, the semipermanent North Pacific High is located in the general area between Hawaii and Alaska. Flow from this High to the California Low results in onshore surface winds along most of the Pacific coast. This seasonal flow, called the Pacific coast monsoon, begins in spring and lasts until fall. The sea breeze is superimposed on the monsoon circulation. During the day, air from the ocean moves inland, rises as it is heated, mixes with the upper winds, and is replaced on the seaward side by gradually settling air from the general circulation.

Since the monsoon flows onshore both day and night, it tends to weaken, or reduce to a negligible amount, the night land breeze. However, this opposition of forces also slows down the onshore monsoon at night. During the day, the sea breeze, assisted by the monsoon, brings in a fresh surge of marine air. Because of this assistance, the marine layer is thicker, and moves farther inland, than does the sea breeze in the East.

The Pacific sea breeze brings relatively cool, moist marine air to the coastal areas. The passage of the leading edge of this air--the sea-breeze front--is marked by a wind shift and an increase in wind speed. Often it is accompanied by fog or low stratus clouds, particularly in the morning hours. Within the first few miles inland, however, the marine air is subjected to heating as it passes over the warmer land. If the marine-air layer is shallower than normal, this air may soon become almost as warm as the air it is replacing. The strong temperature contrasts then remain near the coast while the warmed sea breeze may penetrate many miles beyond.

Thus the effect of the sea breeze on fire behavior can vary considerably. Where the marine air is not modified appreciably, its lower temperatures and higher humidities produce less dangerous fire weather. Where the marine air is modified extensively by heating, the temperature and humidity changes with the sea-breeze front become negligible, while the shifting wind direction and increase in wind speed and gustiness can be a serious detriment to fire control.

Because of surface friction, the sea breeze often moves inland more rapidly at the top of the marine layer than at the surface. Instability and convective mixing caused by surface warming then tend to bring the sea breeze aloft down to the surface, so that the sea-breeze front appears to progress on the surface in jumps or surges. The motion is somewhat analogous to that of the forward portions of the endless metal tracks on a moving tractor.

The Pacific sea breeze is characterized by considerable thermal turbulence and may extend inland 30 to 40 miles or more from the water under favorable conditions. The depth of the sea breeze is usually around 1,200 to 1,500 feet, but sometimes reaches 3,000 feet or more. Its intensity will vary with the water-land temperature contrast, but usually its speed is around 10 to 15 miles per hour.

Mountains along the Pacific coastline act as barriers to the free flow of surface air between the water and the land. On seaward-facing slopes the sea breeze may combine with upslope winds during the daytime, thus transporting modified marine air to the higher elevations in the coastal mountains.

River systems and other deep passes that penetrate the coast ranges provide the principal inland sea-breeze flow routes. The flow of cool, moist air is sufficient to carry tremendous amounts of marine air inland, helping to maintain inland summer humidities at moderate levels in the areas opposite the passes. Here, the sea breeze joins with afternoon upvalley and upcanyon winds, resulting in a cooler, relatively strong flow. In broad valleys, this flow takes on the usual sea-breeze characteristics, but in narrow canyons and gorges it may be strong and very gusty as a result of both mechanical and thermal turbulence.

The coastal mountains similarly cut off major flow from the land to the sea at night. Downslope winds on the ocean-facing slopes join with a feeble land breeze from the coastal strip at night, but again, the outflowing river systems provide the principal flow routes. The downvalley and downcanyon flow is, like the normal land breeze, a relatively shallow and low-speed wind system.

Small-scale diurnal circulations similar in principle to land and sea breezes occur along the shores of inland waters. Lake breezes can appear along the shores of lakes or other bodies of water large enough to establish a sufficient air temperature gradient. The lake breeze is common in summer, for example, along the shores of the Great Lakes. On a summer afternoon it is not unusual for most shore stations to experience onshore winds.

Winds in mountain topography are extremely complex. Part of the time, the general winds associated with larger scale pressure systems dominate the surface layer. But when larger scale pressure systems weaken, the general winds lessen. Then, in the presence of strong daytime heating or nighttime cooling, convective winds of local origin become important features of mountain weather. These conditions are typical of clear summer weather in which there is a large diurnal range of surface air temperatures.

General and convective winds may displace, reinforce, or oppose each other. Their relationship to each other can change quickly, often with surprising rapidity. Variations between different terrain features--sometimes separated only by yards--are often noted. The convective activity may dominate the observed surface wind in one instance, and in another it may permit the speed and direction of winds aloft to dominate the surface flow through the mixing process.

The interactions between airflow of different origins, local pressure gradients caused by non-uniform heating of mountain slopes, and the exceedingly complex physical shapes of mountain systems combine to prevent the rigid application of rules of thumb to convective winds in mountain areas. Every local situation must be interpreted in terms of its unique qualities. Wind behavior described in this section is considered typical, but it is subject to interruption or change at virtually any time or place.

Differences in air heating over mountain slopes, canyon bottoms, valleys, and adjacent plains result in several different but related wind systems. These systems combine in most instances and operate together. Their common denominator is upvalley, upcanyon, upslope flow in the daytime and downflow at night. They result from horizontal pressure differences, local changes in stability that aid vertical motion, or from a combination of the two.

• Slope winds
• Valley winds
• Effects of Orientation and Vegetation on Slope and Valley Winds
• Interaction of Slope and Valley Winds With General Winds

Upslope Winds

Slope winds are local diurnal winds present on all sloping surfaces. They flow upslope during the day as the result of surface heating, and downslope at night because of surface cooling. Slope winds are produced by the local pressure gradient caused by the difference in temperature between air near the slope and air at the same elevation away from the slope.

During the daytime the warm air sheath next to the slope serves as a natural chimney and provides a path of least resistance for the upward flow of warm air. Ravines or draws facing the sun are particularly effective chimneys because of the large area of heated surface and steeper slopes; winds are frequently stronger here than on intervening spur ridges or uniform slopes. Upslope winds are quite shallow, but their depth increases from the lower portion of the slope to the upper portion. Turbulence and depth of the unstable layer increase to the crest of the slope, which is the main exit for the warm air. Here, momentum of the upflowing air, convergence of upslope winds from opposite slopes, and mechanical turbulence combine to make the ridge a very turbulent region where much of the warm air escapes aloft. The crests of higher ridges are also likely to experience the influence of the general wind flow, if that flow is moderate or strong.

An exception to the normal upcanyon, upslope, daytime flow occurs frequently enough on the east slopes of the Pacific Coast Ranges to warrant further discussion; see Downslope Afternoon Winds.

At night the cool air near the surface flows downslope much like water, following the natural drainage ways in the topography. The transition from upslope to downslope wind begins soon after the first slopes go into afternoon shadow and cooling of the surface begins. In individual draws and on slopes going into shadow, the transition period consists of (1) dying of the upslope wind, (2) a period of relative calm, and then (3) gentle laminar flow downslope. Downslope winds are very shallow and of a slower speed than upslope winds. The cooled denser air is stable and the downslope flow, therefore, tends to be laminar.

Downslope winds may be dammed temporarily where there are obstructions to free flow, such as crooked canyons and dense brush or timber. Cool air from slopes accumulates in low spots and overflows them when they are full. The principal force here is gravity. With weak to moderate temperature contrasts, the airflow tends to follow the steepest downward routes through the topography. Strong air temperature contrasts result in relatively higher air speeds. With sufficient momentum, the air tends to flow in a straight path over minor topographic obstructions rather than to separate and flow around them on its downward course.

Cool, dense air accumulates in the bottom of canyons and valleys, creating an inversion which increases in depth and strength during the night hours. Downslope winds from above the inversion continue downward until they reach air of their own density. There they fan out horizontally over the canyon or valley. This may be either near the top of the inversion or some distance below the top.

Theoretically, both upslope and downslope winds may result in a cross-valley circulation. Air cooled along the slopes at night flows downward and may be replaced by air from over the valley bottom. Air flowing upslope in the daytime may be replaced by settling cooler air over the center of the valley. The circulation system may be completed if the upward flowing air, on reaching the upper slopes, has cooled enough adiabatically to flow out over the valley and replace air that has settled. During strong daytime heating, however, cross-valley circulation may be absent. Upflowing air is continually warmed along the slopes. Adiabatic cooling may not be sufficient to offset the warming, and the warmer air is forced aloft above the ridgetops by denser surface air brought in by the upvalley winds.

Downslope Winds

During the forenoon, in the absence of an overriding general wind flow, local winds tend to be upslope and flow up the draws on both the west and east sides of the Coast Ranges. Usually, the flow through gaps and saddles is easterly because of the stronger heating on the east side in the forenoon. The two flows meet in a convergence zone on the west side of the ridge. By midday the flow up the west slopes has increased, most likely because of the sea breeze or a strengthening of the monsoon circulation due to intensification of the thermal trough. The convergence zone has moved eastward across the ridge, and the flow through the gaps has changed to westerly.

Waves form in this westerly flow, which first remain aloft on the lee side of the mountains, and later surface to cause strong downslope winds on the east side. Downslope afternoon winds are commonly three times as strong as the forenoon upslope winds. In some areas, downslope afternoon winds occur nearly every day during the warm season, while in other areas they occur only occasionally. The time of the wind shift from upslope to downslope on the east side may vary from late forenoon to late afternoon, but most frequently it is around noon or early afternoon. On some days, upslope winds redevelop in late afternoon as the mountain waves go aloft. On other days, the downslope afternoon winds diminish and change to the normal nighttime drainage winds.

Valley winds are diurnal winds that flow upvalley by day and downvalley at night. They are the result of local pressure gradients caused by differences in temperature between air in the valley and air at the same elevation over the adjacent plain or larger valley. This temperature difference, and the resulting pressure difference and airflow, reverses from day to night . During the day, the air in the mountain valleys and canyons tends to become warmer than air at the same elevation over adjacent plains or larger valleys.

One reason for the more intensive heating of the mountain valley air is the smaller volume of air in the valley than over the same horizontal surface area of the plain. The rest of the volume is taken up by landmass beneath the slopes. A valley may have only from one-half to three-fourths the volume of air as that above the same horizontal surface area of the plain.

Another reason is the fact that the mountain valley air is somewhat protected by the surrounding ridges from the general wind flow. The valley air is heated by contact with the slopes, and the resulting slope-wind circulation is effective in distributing the heat through the entire mass of valley air. As the valley air becomes warmer and less dense than the air over the plain, a local pressure gradient is established from the plain to the valley, and an upvalley wind begins.

Whereas upslope winds begin within minutes after the sun strikes the slope, the upvalley wind does not start until the whole mass of air within the valley becomes warmed. Usually this is middle or late forenoon, depending largely on the size of the valley. The upvalley wind reaches its maximum speed in early afternoon and continues into the evening. Upvalley wind speeds in larger valleys are ordinarily from 10 to 15 m.p.h. The depth of the upvalley wind over the center of the valley is usually about the same as the average ridge height.

Strong upvalley and upcanyon winds may be quite turbulent because of the unstable air and the roughness of the terrain. Eddies may form at canyon bends and at tributary junctions. Along upper ridges particularly, the flow tends to be quite erratic. Wind speed and direction may change quickly, thus drastically affecting fire behavior.

Slopes along the valley sides begin to cool in late afternoon and, shortly after they come into shadow, cool air starts flowing downslope. Cool air accumulates in the valley bottom as more air from above comes in contact with the slopes and is cooled. Pressure builds up in the valley, causing the upvalley wind to cease. With continued cooling, the surface pressure within the valley becomes higher than the pressure at the same elevation over the plain, and a downvalley flow begins.

The transition from upvalley to downvalley flow takes place in the early night--the time depending on the size of the valley or canyon and on factors favoring cooling and the establishment of a temperature differential. The transition takes place gradually. First, a downslope wind develops along the valley floor, deepens during the early night, and becomes the downvalley wind. The downvalley wind may be thought of as the exodus or release of the dense air pool created by cooling along the slopes. It is somewhat shallower than the upvalley wind, with little or no turbulence because of the stable temperature structure of the air. Its speed is ordinarily somewhat less than the upvalley wind, but there are exceptions in which the downvalley wind may be quite strong. The downvalley wind continues through the night and diminishes after sunrise.

Valley winds and slope winds are not independent. A sloping valley or canyon bottom also has slope winds along its length, although these winds may not be easy to distinguish from valley winds. Proceeding upstream during the daytime, the combined flow continually divides at each tributary inlet into many upravine and upslope components to the ridgetops. As the valley-wind system strengthens during the day, the direction of the upslope wind is affected. The first movement in the morning is directly up the slopes and minor draws to the ridgetop. Then, as the speed of the valley wind picks up, the upslope winds are changed to a more upvalley direction. By the time the valley wind reaches its maximum, the slope winds, on the lower slopes at least, may be completely dominated by the upvalley wind. Along the upper slopes, the direction may continue to be upslope, because the upvalley wind does not always completely fill the valley.

Nighttime downslope winds are similarly affected. When the downvalley wind is fully developed, it dominates the flow along the slopes, particularly the lower portion, so that the observed wind direction is downvalley.

Orientation of the topography is an important factor governing slope-wind and valley-wind strength and diurnal timing. Upslope winds begin as a gentle upflow soon after the sun strikes the slope. Therefore, they begin first on east-facing slopes after daybreak and increase in both intensity and extent as daytime heating continues. South and southwest slopes heat the most and have the strongest upslope winds. South slopes reach their maximum wind speeds soon after midday, and west slopes by about midafternoon. Upslope wind speeds on south slopes may be several times greater than those on the opposite north slopes.

Where slopes with different aspects drain into a common basin, some slopes go into shadow before others and also before the upvalley wind ceases. In many upland basins, the late afternoon upvalley winds are bent in the direction of the first downslope flow. They continue to shift as the downslope flow strengthens and additional slopes become shaded, until a 180-degree change in direction has taken place some time after sunset.

The vegetative cover on slopes will also affect slope winds and, in turn, valley winds. Bare slopes and grassy slopes will heat up more readily than slopes covered with brush or trees (see Effects of Vegetation on Surface Temperatures). Upslope winds will therefore be lighter on the brush- or tree-covered slopes. In fact, on densely forested slopes the upslope wind may move above the treetops, while at the surface there may be a very shallow downslope flow because of the shade provided by the canopy. Downslope winds at night on densely forested slopes are affected by the presence or absence of a dense understory. Where there is an open space between the tree canopy and the surface, the downslope flow will be confined to the trunk region while calm prevails in the canopy region. A forest with a dense understory is an effective barrier to downslope winds. Here, the flow is diverted around dense areas, or confined to stream channels, roadways, or other openings cut through the forest.

Slope wind and valley wind systems are subject to interruption or modification at any time by the general winds or by larger scale convective wind systems such as land and sea breezes.

Midday upslope winds in mountain topography tend to force weak general winds aloft above the ridgetop. The general wind flow goes over the rising currents above the ridge. These rising currents may be effective in producing or modifying waves in the general wind flow. Frequently, the daytime upper winds are felt only on the highest peaks. In this situation, the surface winds, except on the highest peaks, are virtually pure convective winds. Upslope winds dominate the saddles and lower ridges and combine with upvalley winds to determine wind speeds and directions at the lower elevations.

Late afternoon weakening of upslope winds and the onset of downslope flow in the early evening allow the general winds to lower onto the exposed upper slopes and ridgetops. In the Far West, air in the flow aloft from the North Pacific High is subsiding and, therefore, commonly warm and dry. At night, this air may be found at higher levels at least as far inland as the Sierra-Cascade Range. A fire burning to a ridgetop under the influence of upslope afternoon winds may flare up, and its spread may be strongly affected as it comes under the influence of the general wind flow. Similar phenomena may occur in mountainous country elsewhere.

Valley winds are affected by the general wind flow according to their relative strengths, directions, and temperatures. The degree of interaction also varies from day to night. The general wind has its maximum effect on valley winds during the daytime when a strong general wind blows parallel to the valley. If the general wind is blowing in the direction of the upvalley wind and the air is relatively unstable, the influence of the general wind will be felt down to the valley floor. The resulting surface wind will be a combination of the general wind and the upvalley wind. When the general wind blows in the direction opposite to the upvalley wind, it extends its influence some distance down into the valley and the observed surface wind will be the resultant of the upvalley and general winds.

General winds blowing at right angles to the axis of a valley during the daytime have much less influence on the valley wind pattern than those blowing along the valley. The ridges tend to shield the valley circulation from the effects of the general wind.

The relative coldness or density of air being brought in by the general winds is an important factor. Relatively warm air will continue to flow aloft without dropping into valleys and canyons and disturbing the convective wind systems. But cold, relatively dense air combined with strong general wind flow tends to follow the surface of the topography, scouring out valleys and canyons and completely erasing the valley wind systems. Such effects are common in cold air following the passage of a cold front, and in deep layers of cold marine air along the Pacific coast. In these situations the general wind flow is dominant.

These effects are most pronounced when the general wind flow is parallel to the axis of the valley. Strong winds blowing across narrow valleys and canyons may not be able to drop down into them since momentum may carry the airflow across too quickly. Then, too, there are in-between situations where the general wind flow only partly disturbs the valley wind systems. General winds warm adiabatically as they descend the slopes on the windward side of a valley. If the descending air reaches a temperature equal to that of the valley air, it will leave the slope and cross the valley. The cooler the air flowing in with the general wind, the farther it will descend into the valley.

General winds at night usually have much less effect on valley wind systems than during the daytime. Ordinarily a nighttime inversion forms in the valleys, and this effectively shields the downvalley wind from the general wind flow. Again, there are important exceptions that must be considered.

If the air being brought in by the general wind flow is relatively cold and the direction is appropriate, the general wind can combine with downslope and downvalley winds and produce fairly strong surface winds, particularly during the evening hours. Later during the night, however, further cooling will usually establish a surface inversion and the general wind influence will be lifted to the top of the inversion.

Another important exception is the action of lee-side mountain waves. When mountain waves extend down to the surface they will completely obscure valley wind systems. In foehn wind

Whirlwinds

Whirlwinds or dust devils are one of the most common indications of intense local heating. They occur on hot days over dry terrain when skies are clear and general winds are light.

Under intense heating, air near the ground often acquires a lapse rate of 0.2° F per 10 feet which is about 31/2 times the dry-adiabatic rate. The instability is then so extreme that overturning can occur within the layer even in calm air. The superheated air rises in columns or chimneys, establishing strong convective circulations, and drawing in hot air from the surface layer. An upward-spiraling motion usually develops. The spiral is analogous to the whirlpool effect nearly always observed in water draining from a wash basin. The flow becomes spiral because the horizontal flow toward the base is almost invariably off balance.

The lapse rate mentioned in the preceding paragraph is called the autoconvective lapse rate. Greater instability than this may create updrafts spontaneously, but usually a triggering action initiates the updraft. Updrafts can also begin if the layer acquires only a superadiabatic lapse rate; that is, a lapse rate less than the autoconvective but greater than the dry-adiabatic rate. However, with superadiabatic lapse rates, quiet surface air actually remains in vertical equilibrium, and becomes buoyant only if it is lifted. In this case, some triggering action must provide the initial impulse upward. One common triggering action is the upward deflection of the surface wind by an obstacle.

It is probable that nearly all updrafts have some whirling motion, but usually this is weak and invisible. The stronger the updraft, the stronger the whirl, because a larger volume of air is drawn into the vortex. The whirling motion intensifies as the air flows toward the center, much the same as the whirling of an ice skater increases as he moves his arms from an extended position to near his body. The whirl becomes visible if the updraft becomes strong enough to pick up sand, dust, or other debris. The direction of rotation is accidental, depending on the triggering action. It may be either clockwise or counterclockwise.

Whirlwinds may remain stationary or move with the surface wind. If the triggering action is produced by a stationary object, the whirlwind usually remains adjacent to the object. If it does break away, it may die out and another develops over the object. Those whirlwinds that move show a tendency to move toward higher ground. Some whirlwinds last only a few seconds, but many last several minutes and a few have continued for several hours.

The sizes of whirlwinds vary considerably. Diameters range from 10 to over 100 feet, and heights range from 10 feet to 3,000 or 4,000 feet in extreme cases. Wind speeds in the whirlwinds are often more than 20 m.p.h. and in some cases have exceeded 50 m.p.h. Upward currents may be as high as 25 to 30 m.p.h. and can pick up fair-sized debris.

Whirlwinds are common in an area that has just burned over. The blackened ashes and charred materials are good absorbers of heat from the sun, and hotspots remaining in the fire area may also heat the air. A whirlwind sometimes rejuvenates an apparently dead fire, picks up burning embers, and spreads the fire to new fuels. Whirlwinds that develop during fires, an exceedingly dangerous situation, are called firewhirls.

The heat generated by fires produces extreme instability in the lower air and may cause violent firewhirls. Such firewhirls have been known to twist off trees more than 3 feet in diameter. They can pick up large burning embers, carry them aloft, and then spew them out far across the fireline and cause numerous spot fires. At times, the firewhirls move bodily out of the main fire area, but as soon as they do the flame dies out and they become ordinary whirlwinds moving across the landscape.

Firewhirls occur most frequently where heavy concentrations of fuels are burning and a large amount of heat is being generated in a small area. Mechanical forces are often present which serve as triggering mechanisms to start the whirl. A favored area for firewhirls is the lee side of a ridge where the heated air from the fire is sheltered from the general winds. Mechanical eddies produced as the wind blows across the ridge can serve as the triggering mechanism to initiate the whirl. The wind may add to the instability by bringing in cool air at higher levels over the fire-heated air on the lee side. Air streams of unequal speeds or from different directions in adjacent areas can mechanically set off firewhirls in fire-heated air. Firewhirls have also been observed in relatively flat terrain. In these cases the whirls seem to start when a critical level of energy output has been reached by a portion of the fire.

Special winds associated with cumulus cloud growth and thunderstorm development are true convective winds. These winds are

1. the updrafts predominating in and beneath growing cumulus clouds,

2. downdrafts in the later stages of full thunderstorm development, and

3. the cold air outflow which sometimes develops squall characteristics.

There are always strong updrafts within growing cumulus clouds, sometimes 30 m.p.h. or more even if the cumulus does not develop into a thunderstorm. Ordinarily, the air feeding into the cloud base is drawn both from heated air near the surface and from air surrounding the updraft. The indraft to the cloud base may not be felt very far below or away from the cloud cell. A cell that forms over a peak or ridge, however, may actually increase the speed of upslope winds that initiated the cloud formation. A cumulus cloud formed elsewhere that drifts over a peak or ridge also may increase the upslope winds while the cloud grows with renewed vigor. With continued drift, the cloud may draw the ridgetop convection with it for a considerable distance before separating.

If a cumulus cloud develops into a mature thunderstorm, falling rain within and below the cloud drags air with it and initiates a downdraft. Downward-flowing air, which remains saturated by the evaporation of raindrops, is ordinarily warmed at the moist-adiabatic rate. But air being dragged downward in the initial stages of a thunderstorm downdraft is warmed at a lesser rate because of entrainment of surrounding cooler air and the presence of cold raindrops or ice crystals. If this air is dragged downward to a point where it is colder than the surrounding air, it may cascade to the ground as a strong downdraft. In level terrain this becomes a surface wind guided by the direction of the general wind and favorable airflow channels. This is known as the first gust.

In mountainous terrain the thunderstorm downdraft tends to continue its downward path into the principal drainage ways. Speeds of 20 or 30 m.p.h. are common, and speeds of 60 to 75 m.p.h. have been measured. If it is dense enough, the air has sufficient momentum to traverse at least short adverse slopes in its downward plunge. The high speeds and surface roughness cause these winds to be extremely gusty. They are stronger when the air mass is hot, as in the late afternoon, than during the night or forenoon. Although they strike suddenly and violently, downdraft winds are of short duration.

Although downdraft winds are a common characteristic of thunderstorms, it is not necessary for developing cumulus clouds to reach the thunderstorm stage for downdrafts to occur. Downdrafts can develop on hot days from towering cumulus clouds producing only high-level precipitation.

Squall winds often precede or accompany thunderstorms in the mountainous West. These storms often cool sizeable masses of air covering an area of a hundred or several hundred square miles. Occurring as they do in the warm summer months, these cool air masses are in strong temperature contrast with their surroundings. As this air spreads out and settles to lower levels, the leading edge--a front--is accompanied by squall winds. These are strong and gusty; they begin and end quickly. They behave much like wind in squall lines ahead of cold fronts, but are on a smaller geographic scale. However, they may travel out many miles beyond the original storm area.

When air stagnates in a region where surface characteristics are uniform, it acquires those characteristics and becomes an air mass. Warm, moist air masses are formed over tropical waters; cold, moist air masses over the northern oceans; cold, dry air masses over the northern continent; and warm, dry air masses over arid regions. The weather within an air mass--whether cool or warm, humid or dry, clear or cloudy--depends on the temperature and humidity structure of the air mass. Within horizontal layers, the temperature and humidity properties of an air mass are fairly uniform. The depth of the region in which this horizontal uniformity exists may vary from a few thousand feet in cold, winter air masses to several miles in warm, tropical air masses. Weather within an air mass will vary locally from day to day due to heating, cooling, precipitation, and other processes. These variations, however, usually follow a sequence that may be quite unlike the weather events in an adjacent air mass.

Air masses have characteristic weather in their source regions. But, as air masses leave their source regions, they are modified according to the surface over which they travel, and the air-mass weather changes. The day-to-day fire weather in a given area depends, to a large extent, on either the character of the prevailing air mass, or the interaction of two or more air masses. In frontal zones, where differing air masses meet, considerable weather is concentrated. Cloudiness, precipitation, and strong and shifting winds are characteristic of frontal passages; but, occasionally, frontal passages are dry and adversely affect fire behavior.

• Formation and Modification of Air Masses
• Air-Mass Weather
• Fronts
  • Warm Fronts
  • Cold Fronts
  • Stationary Fronts

Air Mass Formation

The region where an air mass acquires its characteristic properties of temperature and moisture is called its source region. Ocean areas, snow- or ice-covered land areas, and wide areas are common source regions. Those areas producing air masses which enter the fire-occurrence regions of North America are:

1. The tropical Atlantic, Caribbean, Gulf Mexico and the tropical Pacific, which are uniformly warm and moist.
2. The Northern Pacific and Northern Atlantic, which are uniformly cool and moist.
3. Interior Alaska, Northern Canada, and the Arctic, which are uniformly cold and dry during the winter months.
4. Northern Mexico and Southwestern States, which are usually hot and dry during the summer months.

The time required for a body of air to come to approximate equilibrium with the surface over which it is resting may vary from a few days to 10 days or 2 weeks, depending largely on whether the body of air is initially colder or warmer than the temperature of its source region. If the air is colder, it is heated from below. Convective currents are produced, which carry the heat and moisture aloft and rapidly modify the air to a considerable height.

On the other hand, if the air is initially warmer than the surface, it is cooled from below. This cooling stabilizes the air and cuts off convection. Cooling of the air above the surface must take place by conduction and radiation, and these are slow processes. Thus, a longer time--up to 2 weeks--is required for the development of cold air masses, and even then, these air masses are only a few thousand feet thick.

Air masses that form over a source region vary in temperature and moisture from season to season, as does the source region. This is particularly true of continental source regions. High-latitude continental source regions are much colder and drier in the winter than in the summer, and tropical continental source regions are much hotter and drier in summer than in winter.

Classification of Air Masses

Air masses are classified according to their source region. Several systems of classification have been proposed, but we will consider only the simplest. Air masses originating in high latitudes are called polar (P), and those originating in tropical regions are called tropical (T). Air masses are further classified according to the underlying surface in the source region as maritime for water and continental for land. The "m" for maritime or "c" for continental precedes the P or T. Thus, the four basic types of air masses, are designated as: mP, mT, cP, and cT, according to their source region. It is natural that air stagnating for some time in a polar region will become cold, or in a tropical region will become warm. And air spending sometime over water becomes moist, at least in the lower layers, while air over land becomes dry.

Air Mass Modification

As an air mass leaves its source region in response to broadscale atmospheric motions, it may be colder or warmer than the surface it passes over. It is then further classified by the addition of k for colder or w for warmer to its classification symbol. The k-type air mass will be warmed from below and will become unstable in the lower layers. A w-type air mass will be cooled from below, will become stable, and will be modified slowly, and only in the lower few thousand feet.

Air-mass properties begin changing as soon as the air mass leaves its source region. The amount of modification depends upon the speed with which the air mass travels, the type of surface over which it moves, and the temperature difference between the air mass and the underlying surface.

Air masses are modified in several ways. For the most part, these are processes which we have already considered in detail. Several of the processes usually take place concurrently:

• An air mass is heated from below if it passes over a warmer surface (previously warmed by the sun) or if the surface beneath a slow-moving air mass is being currently warmed by the sun. Such modification is rapid because of the resulting instability and convection.
• An air mass is cooled from below if it passes over a colder surface, or -if the surface is cooled by radiation. This increases the stability of the lower layers, and further modification becomes a slow process.
• Moisture may be added to an air mass by:
  • Evaporation from water surfaces, moist ground, and falling rain;
  • sublimation from ice or snow surfaces and falling snow or hail; and
  • transpiration from vegetation. Of these, sublimation is a relatively slow process by comparison.
• Moisture may be removed from an air mass by condensation and precipitation.
• Finally, air-mass properties may be changed by turbulent mixing, by sinking, or by lifting.

After moving a considerable distance from its source region, particularly after entering a source region of another type, an air mass may lose its original distinctive characteristics entirely and acquire those of another air-mass type. Thus, a continental polar air mass moving out over the Gulf of Mexico takes on the characteristics of a maritime tropical air mass. Or a maritime polar air mass, after crossing the Rocky Mountains, may assume the characteristics of a continental polar air mass.

There are many differences in air masses and in the weather associated with them. Even within one air-mass type, there will be considerable variation, depending on the season, the length of time that an air mass has remained over its source region, and the path it has followed after leaving that region. We will discuss only the most common characteristics on only the more distinct types of air masses:

continental polar- winter maritime polar- winter maritime tropical winter

continental polar- summer maritime polar- summer maritime tropical- summer continental tropical- summer

Following these descriptions of common air masses, consider a few general principles to help us understand the many variations in individual air masses.

Continental Polar--Winter

Continental polar air masses originate in the snow-covered interior of Canada, Alaska, and the Arctic in the colder months. Lower layers of the air become quite cold, dry, and stable. Much moisture from the air is condensed onto the snow surface. These air masses are high-pressure areas, and there is little cloudiness due to the lack of moisture and to the stability of the air mass.

These are the coldest wintertime air masses, and cause severe cold waves when moving southward through Canada and into the United States. Upon moving southward or southeastward over warmer surfaces, cP air masses change to cPk. The lower layers become unstable and turbulent. If a part of the air mass moves over the Great Lakes, it picks up moisture as well as heat and may produce cloudiness and snow flurries or rain showers on lee side of the Lakes, and again on the windward side of the Appalachian Mountains. Once across the Appalachians, the air mass is generally clear and slightly warmer.

If a cP air mass moves southward into the Mississippi Valley and then into the Southeast, it will gradually warm up but remain dry. Modification is slow until the air mass passes beyond the snow-covered areas, and then it becomes more rapid. When cP air moves out over the Gulf of Mexico, it is rapidly changed to an mT air mass. The generally clear skies and relatively low humidities associated with cP air masses are responsible for much of the hazardous fire weather in the South and Southeast during the cool months.

The Rocky Mountains effectively prevent most cP air masses from moving into the Far West. But occasionally, a portion of a deep cP air mass does move southward west of the Rockies, and in so doing brings this area its coldest weather. At times, the air is cold enough for snow to fall as far south as southern California.

Maritime Polar--Winter

The North Pacific is the common source region for maritime polar air masses. While in its source region, the air mass is cold and has a lapse rate nearly the same as the moist-adiabatic rate. If the air mass moves into the snow-covered regions of Canada, it gradually changes to a cP air mass. Maritime polar air taking that trajectory usually has had a comparatively short stay over the water. It is quite cold and has high relative humidity, but moisture content in terms of absolute humidity is rather low. However, rain or snow showers usually result as the air is lifted over the coastal mountains.

Maritime polar air masses originating farther south and entering Western United States or Southwestern Canada have had a longer overwater trajectory, are not quite so cold, and have a higher moisture content. On being forced over the Coast Ranges and the Rocky Mountains, an mP air mass loses much of its moisture through precipitation. As the air mass descends on the eastern slopes of the Rocky Mountains, it becomes relatively warm and dry with generally clear skies. If, however, it cannot descend on the lee side of the mountains, and instead continues eastward over a dome of cold cP air, snow may occur.

East of the Rockies, mP air at the surface in winter is comparatively warm and dry, having lost much of its moisture in passing over the mountains. Skies are relatively clear. If this air mass reaches the Gulf of Mexico, it is eventually changed into an mT air mass.

Maritime polar air sometimes stagnates in the Great Basin region of the Western United States in association with a Great Basin High. The outflow from the Great Basin High may give rise to strong, dry foehn winds in a number of the surrounding States.

At times during the winter, mP air is trapped in Pacific coast valleys and may persist for a week or more. Low stratus clouds and fog are produced, making these valleys some of the foggiest places on the continent during the winter.

Although mP air forms over the North Atlantic Ocean, as well as the North Pacific, the trajectory of Atlantic mP air is limited to the northeastern seaboard.

Maritime Tropical--Winter

Most of the maritime tropical air masses affecting temperate North America originate over the Gulf of Mexico or Caribbean Sea. They are warm, have a high moisture content, and a conditionally unstable lapse rate. Maritime tropical air is brought into the southeastern and central portions of the country by the circulation around the western end of the Bermuda High. In moving inland during the winter, mT air is cooled from below by contact with the cooler continent and becomes stabilized in the lower levels. Fog and low stratus clouds usually occur at night and dissipate during the day as this air mass invades the Mississippi Valley and the Great Plains. If mT air is lifted over a cP air mass, or if it moves northeastward and is lifted on the western slopes of the Appalachians, the conditional instability is released and large cumulus clouds, heavy showers, and frequent thunderstorms result.

Maritime tropical air seldom reaches as far as the Canadian border or the New England States at the surface in winter. Nevertheless, it occasionally causes heavy rain or snow in these areas, when mT air encounters a colder cP or mP air mass and is forced to rise up over the denser air. More will be said about this process in the section on fronts.

The tropical Pacific is also a source region for mT air, but Pacific mT seldom enters the continent. When it does, it is usually brought in with a low-pressure system in Northern Mexico or California, where the Pacific mT air can cause heavy rainfall when rapidly forced aloft by the mountains.

Continental Polar--Summer

In summer, even though the source region for cP air masses is farther north than in winter-over Northern Canada and the polar regions-the warmer surface temperatures result in little surface cooling and frequently in actual heating of the air near the ground. The air mass, therefore, may be relatively unstable in the lower layers in contrast to its extreme stability during the winter. Since the air is dry from the surface to high levels, the relative instability rarely produces cloudiness or precipitation.

The general atmospheric circulation is weaker during the summer, and polar outbreaks move more slowly than in winter. As a result, cP air undergoes tremendous changes in passing slowly from its source region to Southern United States. During its southward and southeastward travel, cP air is warmed from below and becomes more unstable.

Continental areas, over which cP air travels, are relatively moist in summer, being largely covered with crops, grass, forests and other vegetation. Transpiration from these plants and evaporation from water bodies and moist soil increase the moisture content of cP air rather rapidly. As the moisture content increases, cloudiness also increases.

The weather associated with cP air as it passes through Canada and enters the United States is generally fair and dry. Frequent intrusions of this air give rise to much of the fire weather in the north-central and northeastern regions from spring, through summer, and into fall.

Occasionally, cP air stagnates in the Southeastern United States and accumulates sufficient moisture to produce showers and isolated thunderstorms, particularly over mountainous areas.

Maritime Polar--Summer

Maritime polar air masses in summer originate in the same general area over the Pacific Ocean as in winter. In summer, however, the ocean is relatively cool compared to the land surfaces. Summer mP air is cooled from below in its source region and becomes stable. Stability in the lower layers prevents moisture from being carried to higher levels. Aloft, this air mass remains very dry, usually even drier than summer cP, and becomes quite warm through the subsidence, which takes place in the Pacific High.

As mP air approaches the Pacific coast, the cold, upwelling waters along the shore cause further cooling, increasing relative humidity, and stimulating the formation of considerable fog or low stratus clouds. Thus, along the Pacific coast, summer mP is characterized by a cool, humid marine layer from 1,000-2,000 feet thick, often with fog or low stratus clouds, a strong inversion capping the marine layer, and warm, dry, subsiding air above.

As mP air moves inland from the west coast, the strong daytime heating in interior California, Oregon, Washington, and portions of British Columbia warms the surface layers and lowers the relative humidity. The intense heating and the lifting as mP air crosses the mountains may result in cumulus cloud formation and occasional scattered showers and thunderstorms at high elevations. In descending the eastern slopes of the Rockies, summer mP is heated adiabatically as in winter, and the relative humidity may become quite low at times. When it arrives in the Plains and the Mississippi Valley, it is hardly distinguishable from cP air in the area and results in clear, dry weather. Continuing eastward, it becomes warmer and more unstable, and picks up moisture from the earth and plants. By the time it reaches the Appalachians, it has become unstable and moist enough so that lifting can again produce showers or thunderstorms.

Maritime polar air formed over the colder waters of the North Atlantic in summer occasionally moves southward bringing cool weather and cloudiness to the Atlantic coastal areas.

Maritime Tropical--Summer

Maritime tropical air in its source region over the Gulf of Mexico and the Caribbean in summer has properties similar to those in winter, except that it is conditionally unstable to higher levels, slightly warmer, and moister. In summer, mT air invades central and eastern North America very frequently, sometimes penetrating as far north as Southern Canada, bringing with it the typical heat and oppressive humidity of those tropical source regions.

Daytime heating of the air as it moves inland produces widespread showers and thunderstorms, particularly, during the afternoon and evening. At night, there may be sufficient cooling of the earths surface to bring the temperature of the air near the ground to the dew point and produce fog or stratus clouds. This is dissipated in the early morning by surface heating,

When mT air is lifted, either by crossing mountains or by being forced to rise over cooler mP or cP air, widespread clouds, numerous showers, and intense thunderstorms are produced.

Although some of the summer thunderstorm activity in Northern Mexico and the Southwestern United States is the result of mT air from the tropical Pacific, most of it is associated with mT air from the Gulf of Mexico. This moist air is usually brought in at intermediate levels by easterly and southeasterly flow. Heating and lifting by mountains set off thunderstorms as the air spreads northward, along the Sierra-Cascade range, occasionally extending as far as northern Idaho, western Montana, and Southern Canada. Some thunderstorm activity develops as mT air spreads northwestward from the Gulf and is lifted along the eastern slopes of the Rocky Mountains.

On rare occasions, mT air originating in the tropical Pacific spreads northward over Northwestern Mexico and California with thunderstorm activity. Usually this is residual mT air from a dying tropical storm.

Continental Tropical--Summer

The only source regions for continental tropical air in North America are Mexico and the Southwestern United States. This air is hot, dry, and unstable, and causes droughts and heat waves when it persists for any length of time. It is similar to the upper-level subsiding air in the Pacific High, and may actually be produced by subsidence from aloft.

In summer, cT air sometimes spreads eastward and northward to cover portions of the Central or Western United States. Because of its heat and dryness, it has a desiccating effect on wildland fuels, setting the stage for serious fire-weather conditions.

Variations in Air-Mass Weather

We have considered the usual characteristics of the principal air masses in winter and in summer. We must realize, however, that there are many variations in individual air masses--variations from day to night, and seasonal variations other than just in winter and summer. We will consider a few general principles to help us understand these variations.

1. If the surface over which an air mass is located is warmer than the air mass, the lower layers will be heated. This results in increased instability, convective mixing and turbulence, and a lowering of surface relative humidity. If sufficient moisture is present, cumulus clouds and possible showers may be formed. The increased mixing generally results in good visibility.
2. If the surface is colder than the air mass, the lower layers are gradually cooled. This increases the stability and retards convective mixing and turbulence. Water vapor and atmospheric impurities tend to be concentrated in the lower layers, and visibility is decreased. With sufficient moisture, fog and low stratus clouds will form.
3. As a rule, air masses over land and away from their source region tend to be cooler than the surface during the day and warmer than the surface at night. Thus, the weather characteristics change accordingly from day to night.
4. In the spring, land surfaces away from source regions warm faster than the water or snow-covered surfaces at source regions. This leads to increased instability in the lower layers as air masses leave their source region, and causes considerable thunderstorm, activity, hail, and, sometimes, tornadoes.
5. During the summer, there is the least temperature difference between polar and tropical regions. The general circulation is weaker so that air masses move more slowly, and spending more time in transit, are thus more subject to modification. The belt of westerlies is farther north than in winter. As a result, tropical air masses penetrate far to the north, but polar air masses are blocked at high latitudes and do not penetrate far southward.
6. As the earths surface begins to cool in the fall, air masses tend to be more stable in the lower layers, and thunderstorm activity is reduced. As fall progresses and winter approaches, stable cold air near the surface becomes deeper and more persistent, encouraging the formation of fog or low stratus clouds.
7. During the winter, cold polar air masses move at a faster rate and penetrate far southward. The temperature contrast between polar and tropical regions increases, as does the speed of the general circulation.

Some of the weather conditions most adverse to fire control, such as strong, gusty winds, turbulence, and lightning storms, occur in frontal zones. Sometimes there is insufficient moisture in the warm air mass, or inadequate lifting of this mass, so that no precipitation occurs with the front. Strong, gusty, and shifting winds are typical of a dry frontal zone, adding greatly to the difficulty of fire control.

Types of fronts are distinguished by the way they move relative to the air masses involved. If a front is moving so that cold air is replacing warm air, it is a cold front. If the warm air is advancing and replacing cold air ahead, the front is a warm front. If a front is not moving, it is a stationary front. Cold fronts are indicated on weather maps by pointed cusps, and warm fronts by semicircles, on the side toward which they are moving. A stationary front is indicated by a combination of both.

In a frontal zone, the warmer air mass, being lighter, will be forced over the colder air mass. The rotation of the earth deflects the movement of both the cold and the warm air masses as one tries to overrun or underride the other, and prevents the formation of a horizontal discontinuity surface. Instead, the frontal surface slopes up over the colder air. The slope varies from about 1/50 to 1/300. A 1/50 slope means that for every 50 miles horizontally, the front is 1 mile higher in the vertical. The amount of slope is dependent upon the temperature contrast between the two air masses, the difference in wind speed across the front, and the relative movements of the air masses involved; that is, whether cold air is replacing warm air at the surface or warm air is replacing cold air. On a surface weather map, only the intersection of the frontal surface with the earth is indicated. The contrast between the air masses is strongest near the earths surface, and decreases upward in the atmosphere.

The central portions of air masses are usually associated with areas of high pressure, but fronts are formed in troughs of low pressure. From a position on a front, we find that the pressure rises both toward the warmer air and toward the colder air. Because the gradient wind in the Northern Hemisphere always blows with high pressure on the right, as one faces downstream, this means that the wind blows in one direction in the cold air and a different direction in the warm air. At a given location the wind shifts in a clockwise direction as a front passes--for example, from southeast to southwest or from southwest to northwest.

The wind-shift line and pressure trough line provide good clues to the weatherman for the location of fronts, but there are other indications to consider. A temperature discontinuity exists across a front. As a rule, the greater and more abrupt the temperature contrast, the more intense the front. Weak fronts are characterized by gradual and minor changes in temperature. The moisture contrast between air masses on different sides of a front may be indicated by the dew-point temperatures. Usually the cold air mass will be drier than the warm air mass. Other indications of front location are cloud types, pressure changes, and visibility changes.

The leading edge of an advancing cold air mass is a cold front. It forms a wedge, which pushes under a warm air mass forcing the warm air to rise. Because of surface friction, the lowest layers of the cold air are slowed down. This increases the steepness of the frontal surface and causes a cold front to have a blunted appearance when viewed in cross-section. The slopes of cold fronts usually vary from 1/50 to 1/150.

There are wide variations in the orientation and speed of cold fronts. Usually, they are oriented in a northeast-southwest direction, and they move to the east and southeast, at speeds varying from about 10 to 40 m.p.h. and faster in the winter.

As a cold front approaches, the southerly winds increase in the warm air ahead of the front. Clouds appear in the direction from which the front is approaching. The barometric pressure usually falls, reaches its lowest point as the front passes, then rises sharply. Winds become strong and gusty and shift sharply to westerly or northwesterly as the cold front passes. Temperature and dew point are lower after the cold front passes. In frontal zones with precipitation, the heaviest precipitation usually occurs with the passage of the front. Then it may end quickly and be followed by clearing weather.

There are many exceptions to the foregoing general pattern of cold-front passages. The severity of the weather associated with cold fronts depends upon the moisture and stability of the warm air, the steepness of the front, and the speed of the front. Since cold fronts are usually steeper and move faster than warm fronts, the accompanying band of weather is narrower, more severe, and usually of shorter duration than with warm fronts.

With slow-moving cold fronts and stable warm air, rain clouds of the stratus type form in a wide band over the frontal surface and extend for some distance behind the front. If the warm air is moist and conditionally unstable, thunderstorms may form, with the heaviest rainfall near the frontal zone and immediately following. If the warm air is fairly dry and the temperature contrast across the front is small, there may be little or no precipitation and few or no clouds.

With rapidly moving cold fronts, the weather is more severe and occupies a narrower band. The disturbance is also of shorter duration than that caused by a slow-moving front. If the warm air is relatively stable, overcast skies and precipitation may occur for some distance ahead of the front, and the heaviest precipitation may occur ahead of the surface cold front. If the warm air is moist and conditionally unstable, scattered showers and thunderstorms form just ahead of the cold front. The weather usually clears rapidly behind a fast-moving cold front, with colder temperatures and gusty, turbulent surface winds following the frontal passage.

Under some conditions, a line of showers and thunderstorms is formed from 50 to 300 miles ahead of, and roughly parallel to, a cold front. This is called a squall line. The weather associated with squall lines is often more severe than that associated with the subsequent cold front. After the passage of the squall line, the temperature, wind, and pressure usually revert to conditions similar to those present before the squall line approached. Occasionally, the showers and thunderstorms are scattered along the squall line so that some areas experience strong, gusty winds without any precipitation.

Dry cold fronts often cause very severe fire weather in many sections. Dry cold-front passages may occur in any region, but they are a major problem in the Southeast. Cold fronts tend to be drier farther away from the low-pressure center with which they are associated. Thus, a cold front associated with a Low passing eastward across Southern Canada or the Northern States may be very dry as it passes through the Southeast. In addition, the polar air mass following the cold front may become quite unstable because of surface heating by the time it reaches the Southeast.

The combination of strong, gusty winds and dry, unstable air creates serious fire weather. The second of two cold fronts passing through the Southeast in rapid succession also tends to be dry. The warm air mass ahead of the first cold front may be moist and produce precipitation, but the air mass between the first and second fronts usually will not have had time to acquire much moisture. Therefore, the second cold-front passage may be dry and will be the more serious from the fire-control standpoint.

The dry, trailing ends of cold fronts cause serious fire weather wherever they occur. Along the Pacific coast, the winds behind such cold fronts are, at times, from a northeasterly direction. This offshore direction means that the air flows from high elevations to low elevations and has foehn characteristics. The strong, shifting, gusty winds of the cold-front passage combine with the dry foehn wind to the rear of the front to produce a short-lived but extremely critical fire-weather condition.

The leading edge of an advancing warm air mass is called a warm front. The warm air is overtaking and replacing the cold air, but at the same time sliding up over the wedge of cold air. Warm fronts are flatter than cold fronts, having slopes ranging from 1/100 to 1/300. Because of this flatness, cloudiness and precipitation extend over a broad area ahead of the front, providing, of course, that there is sufficient moisture in the warm air.

Warm fronts are less distinct than cold fronts and more difficult to locate on weather maps. This is particularly true in rough terrain where high-elevation areas may extend up into the warm air before the warm front has been felt at lower elevation stations.

The first indication of the approach of warm, moist air in the upper levels ahead of the surface warm front may be very high, thin, cirrostratus clouds, which give the sky a milky appearance. These are followed by middle-level clouds, which darken and thicken as precipitation begins. This sequence may be interrupted by short clearing periods, but the appearance of successively lower cloud types indicates the steady approach of the warm front. Rains may precede the arrival of the surface warm front by as much as 300 miles. Rain falling through the cold air raises the humidity to the saturation level and causes the formation of low stratus clouds.

If the warm air above the warm front is moist and stable, the clouds that form are of the stratus type. The sequence is cirrus, cirrostratus, altostratus, and nimbostratus. Precipitation is a steady type and increases gradually with the approach of the surface front.

If the warm air is moist and conditionally unstable, altocumulus and cumulonimbus clouds and, frequently, thunderstorms will be embedded in the cloud masses that normally accompany a warm front.

The rate of movement of warm fronts is about half that of cold fronts. Winds are usually not as strong or gusty with the approach of fronts as with cold fronts. The shift in wind is generally from an easterly to a southerly direction as a warm front passes. After it passes, temperatures rise, precipitation usually stops, and clouds diminish or vanish completely.

From the standpoint of fire weather, warm fronts associated with moist air are a real benefit. The accompanying precipitation is widespread and long-lasting, and usually is sufficient to thoroughly moisten forest fuels, reducing the fire danger.

When the forces acting on two adjacent air masses are such that the frontal zone shows little movement, the front is called a stationary front. Surface winds on either side of the front tend to blow parallel to the front, but in opposite directions. Weather conditions occurring with a stationary front are variable; usually they are similar to those found with a warm front, though less intense. If the air is dry, there may be little cloudiness or precipitation. If the air is moist, there may be continuous precipitation with stable, warm air, or showers and thunderstorms with conditionally unstable, warm air. The precipitation area is likely to be broader than that associated with a cold front, but not as extensive as with a warm front.

Stationary fronts may quickly change back to moving fronts as a slight imbalance of forces acting on the air masses develops. A stationary front may oscillate back and forth, causing changing winds and weather conditions at a given location. It may become a cold or warm front, or a frontal wave may develop.

A frontal surface is similar to a water surface. A disturbance such as wind can cause the formation of waves on the water. If the wave moves toward the shoreline, it grows until it becomes top-heavy and breaks. Similarly, along frontal surfaces in the atmosphere a disturbance may form a wave. This disturbance may be a topographic irregularity, the influence of an upper-level trough, or a change in the wind field cause by local convection. Waves usually form on stationary fronts or slow-moving cold fronts, where winds on the two sides of the front are blowing parallel to the front with a strong shearing motion.

When a section of a front is disturbed, the warm air begins to flow up over and displace some of the cold air. Cold air to the rear of the disturbance displaces some of the warm air. Thus, one section of the front begins to act like a warm front, and the adjacent section like a cold front. This deformation is called a frontal wave.

The pressure at the peak of the frontal wave falls, and a low-pressure center with a counterclockwise (cyclonic) circulation is formed. If the pressure continues to fall, the wave may develop into a major cyclonic system. The Low and its frontal wave generally move in the direction of the wind flow in the warm air, which is usually toward the east or northeast.

As the system moves, the cold front moves faster than the warm front and eventually overtakes the warm front. The warm air is forced aloft between the cold air behind the cold front and the retreating cold air ahead of the warm front. The resulting combined front is called an occlusion or occluded front. This is the time of maximum intensity of the wave cyclone. The pressure becomes quite low in the occluded system with strong winds around the Low. Usually the system is accompanied by widespread cloudiness and precipitation. The heaviest precipitation occurs to the north of the low-pressure center.

As the occlusion continues to grow in length, the cyclonic circulation diminishes in intensity, the low-pressure center begins to fill, and the frontal movement slows down.

There are two types of occluded fronts a warm-front type and a cold-front type depending on whether the surface air ahead of the occlusion is warmer or colder than the air to the rear.

The cold-front type is predominant over most of the continent, especially the central and eastern regions. The weather and winds with the passage of a cold-front occlusion are similar to those with a cold front. Ahead of the occlusion, the weather and cloud sequence is much like that associated with warm fronts.

Most warm-front occlusions are found along the west coast. The air mass to the rear is warmer than the air mass ahead. Therefore, when the cold front overtakes the warm front, it rides up the warm-front surface and becomes an upper cold front.

The weather associated with a warm-front occlusion has characteristics of both warm-front and cold-front weather. The sequence of clouds and weather ahead of the occlusion is similar to that of a warm front. Cold-front weather occurs near the upper cold front. With moist and conditionally unstable air, thunderstorms may occur. At the surface, the passage of a warm-front occlusion is much like that of a warm front. The rainy season in the Pacific Northwest, British Columbia, and southeastern Alaska is dominated by a succession of warm front occlusions that move in from the Pacific.

Another type of upper cold front should be mentioned. Cold fronts approaching the Rocky Mountains from the west are forced to rise and cross over the mountains. Quite frequently in winter, a very cold air mass is located east of the mountains. Then, the cold front does not return to the surface, but rides aloft over the cold air as an upper cold front often accompanied by thundershowers. When such a front meets an mT air mass, and underrides it, a very unstable condition is produced that will result in numerous thunderstorms and, occasionally, tornadoes.

Clouds are visible evidence of atmospheric moisture and atmospheric motion. Those that indicate instability may serve as a warning to the fire-control managers. Some produce precipitation and become an ally to the firefighter. Some clouds develop into full-blown thunderstorms with firestarting potential and often disastrous effects on fire behavior. The amount of precipitation and its seasonal distribution are important factors in controlling the beginning, ending, and severity of local fire seasons. Prolonged periods with lack of clouds and precipitation set the stage for severe burning conditions by increasing the availability of dead fuel.

Clouds consist of minute water droplets, ice crystals, or a mixture of the two in sufficient quantities to make the mass discernible. Air becomes saturated either by the addition of moisture, or, more commonly, by cooling to the dew point. In saturated air, clouds form by the condensation of water vapor, which takes place on fine particles called condensation or sublimation nuclei. Cloud droplets grow to sizes large enough to precipitate by the ice-crystal process, in which water vapor is transferred from evaporating, supercooled liquid droplets to ice crystals where sublimation takes place, or by coalescence of droplets or ice crystals into raindrops or clumps of snowflakes. Precipitation falls in the form of liquid rain or drizzle, freezing rain or drizzle, or frozen snow, sleet, or hail.

Clouds are classified according to their structure as stratus or cumulus, and according to their altitude as high, middle, low clouds, and those with large vertical development. In the last group are cumulonimbus or thunderstorm clouds.

• Saturation Processes
• Condensation, Sublimation, and Precipitation
• Kinds of Precipitation
• Kinds of Clouds

In order for clouds to form and precipitation to develop, the atmosphere must be saturated with moisture. At saturation, the atmospheric vapor pressure is equal to the saturation vapor pressure at the existing temperature and pressure. There are two principal ways in which the atmospheric vapor pressure and saturation vapor pressure attain the same value to produce 100 percent relative humidity, or saturation:

• through the addition of moisture to the air, or more importantly
• through the lowering of air temperature.

As cold air passes over warm water, rapid evaporation takes place, and saturation is quickly reached. Cold continental polar air crossing the warmer Great Lakes in the fall and early winter gathers large amounts of moisture and produces cloudiness and frequently causes rain or snow to the lee of the lakes. Saturation may also be reached when warm rain falls through cold air; for example, beneath a warm front. Rain falling from the warm clouds above the front evaporates in the cold air beneath and forms scud clouds. Contrails made by high-flying aircraft are a type of cloud formed by the addition of moisture from the planes exhaust.

The more important method of reaching saturation, by lowering air temperature, is accomplished in several ways. Warm, moist air may be cooled to its dew point by passing over a cold surface; the cooling takes place near the surface so that, with light wind conditions, fog is formed. If the winds are strong, however, they will cause mixing of the cooled air, and clouds will form several hundred or even a thousand or more feet above the surface.

Lifting of air, and the resultant adiabatic expansion, is the most important cooling method. It produces most of the clouds and precipitation. The lifting may be accomplished by thermal, orographic, frontal action, or convergence.

Thermal Lifting

Local heating will result in thermal lifting. As heated surface air becomes buoyant, it is forced aloft and cools. The air cools at the dry-adiabatic rate of 5.5° F per thousand feet, while the dew point lowers only about 1° F per thousand feet, reflecting the decreasing absolute humidity with expansion. Thus, the temperature and dew point approach each other at the rate of 4.5° F per thousand feet. If the locally heated air contains enough moisture and rises far enough, saturation will be reached and cumulus clouds will form. In fact, a common method of estimating the base of cumulus clouds formed, usually in the summer months, by thermal convection in the lower layers is to divide the difference between the surface air temperature and dew point by 4.5. This gives the approximate height of the cloud base in thousands of feet.

As an example, suppose we begin with heated air at the surface having a temperature of 84° F, wet-bulb temperature of 71° F, dew-point temperature of 66° F, and a relative humidity of 54 percent. If the air rose to an altitude of 4,000 feet, the dry-bulb, wet-bulb, and dew-point temperatures would all have decreased to 62° F. Saturation would have been reached as the humidity would be 100 percent. Continued rising would produce condensation and visible clouds.

Thermal lifting is most pronounced in the warm seasons. It may turn morning stratus clouds into stratocumulus with the possibility of light showers. More frequently, depending on stability, continued heating develops cumuliform clouds that result in heavier showers and thunderstorms. Rainfall associated with thermal lifting is likely to be scattered in geographic extent. In flat country, the greatest convective activity is over the hottest surfaces. In mountain country, it is greatest over the highest peaks and ridges.

Orographic Lifting

Orographic lifting, in which air is forced up the windward side of slopes, hills, and mountain ranges, is an important process in producing clouds and precipitation. As in thermal lifting, the air is cooled by the adiabatic process.

In the West, maritime polar air flowing in from the Pacific Ocean produces winter clouds and precipitation as it is lifted over the mountain ranges. The Coast Ranges, Sierra-Cascades, and Rocky Mountains are the principal mountain systems involved. Lifting in each case occurs on the western slopes, and it is these that receive the heaviest precipitation. The lee slopes and adjacent valleys and plains receive progressively less as the air moves eastward.

Similarly, in the East, maritime tropical air that has moved into the central portion of the United States and Southern Canada is lifted and produces precipitation in the Appalachian Mountains as it progresses eastward. Other air masses, such as continental polar and maritime polar, will also cause precipitation in these mountains if they have acquired sufficient moisture before being lifted.

Frontal Lifting

Frontal lifting, as air is forced up the slope of warm or cold fronts, accounts for much cloudiness and precipitation in all regions in the winter and in many regions during all seasons of the year. East of the Rockies and along the west coast, warm fronts, because of the gradual slope of their frontal surfaces, typically produce steady rains over extensive areas. Cold fronts, with characteristically steeper and faster moving leading surfaces, frequently produce more intense rainfall from cumulonimbus clouds along the front or along a squall line ahead of the front. This rainfall, however, is usually more scattered and of a shorter duration than that produced by a warm front.

Convergence

Convergence is another important method of lifting which produces clouds and precipitation. During convergence, more air moves horizontally into an area that moves out. The excess is forced upward. Since moisture is concentrated in the lower atmospheric levels, convergence, like other lifting mechanisms, carries large quantities of moisture to higher levels. Even when precipitation does not immediately result from this cause, subsequent precipitation triggered by other processes may be much more intense than if convergence had not occurred.

The circulation around a low-pressure system results in convergence. Here, friction deflects the flow toward the center. With more air flowing toward the center than away from it, there is a corresponding upward flow of air. For this reason, low-pressure areas are usually areas of cloudiness and precipitation. On a small scale, convergence occurs during the daytime over mountain peaks and ridges as thermal up-slope winds from opposing sides meet at the top.

We have discussed various methods by which air becomes saturated and condensation and precipitation are produced, but we must remember that in most cases two or more of these methods are acting at the same time. Daytime cumulus clouds over mountains may be produced by heating, orographic lifting, and the convergence of thermal winds all acting together. Nighttime fog and drizzle in maritime tropical (mT) air that moves from the Gulf of Mexico into the Plains regions may be the result of a combination of orographic lifting and nighttime cooling. Frontal lifting may be assisted by orographic lifting in mountain areas, or by convergence in low-pressure areas and troughs.

For condensation or sublimation to occur in the free air, a particle or nucleus must be present for water-vapor molecules to cling to. These fine particles are of two types: condensation nuclei and sublimation nuclei. Condensation nuclei, on which liquid cloud droplets form, consist of salt particles, droplets of sulfuric acid, and combustion products. They are usually abundant in the atmosphere so that cloud droplets form when saturation is reached. Sublimation nuclei, on which ice crystals form, consist of dust, volcanic ash, and other crystalline materials. Because of differences in composition and structure, different nuclei are effective at different below-freezing temperatures. As the temperature decreases, additional nuclei become active in the sublimation process. These nuclei are not as plentiful as condensation nuclei. Even at temperatures well below freezing, there frequently are too few effective nuclei to initiate more than a scattering of ice crystals.

Condensation

The small particles that act as condensation nuclei are usually hygroscopic; that is, they have a chemical affinity for water. They may absorb water well before the humidity reaches saturation, sometimes at humidities as low as 80 percent. Condensation forms first on the larger nuclei and a haze develops which reduces visibility. As the relative humidity increases, these particles take on more water and grow in size while condensation also begins on smaller nuclei. Near saturation, the particles have become large enough to be classed as fog or cloud droplets, averaging 1/2500 inch in diameter, and dense enough so that the mass becomes visible. Rapid cooling of the air, such as in strong upward currents, can produce humidities of over 100 percent--supersaturation--temporarily. Under such conditions, droplets grow rapidly, very small nuclei become active and start to grow, and many thousands of droplets per cubic inch will form. With supersaturation, even nonhygroscopic particles will serve as condensation nuclei, but usually there are sufficient hygroscopic nuclei so that the others do not have a chance.

As condensation proceeds, droplets continue to grow until they reach a maximum size of about 1/100 inch in diameter, the size of small drizzle drops. The condensation process is unable to produce larger droplets for several reasons. As vapor is used up in droplet formation, supersaturation decreases and the cloud approaches an equilibrium state at saturation. Also, as droplets grow, the mass of water vapor changing to liquid becomes large and the resultant latent heat released in the condensation process warms the droplet and decreases the vapor pressure difference between it and the surrounding vapor. Thus the vast majority of clouds do not produce rain. If growth to raindrop size is to take place, one or more of the precipitation processes (ice-crystal and coalescence processes, discussed below) must come into play.

An important phenomenon in the physics of condensation and precipitation is that liquid cloud droplets form and persist at temperatures well below freezing. Although ice melts at ~32° F, water can be cooled much below this before it changes to ice. Liquid cloud droplets can exist at temperatures as low as - 40° F. More commonly in the atmosphere though, cloud droplets remain liquid down to about 15° F. Liquid droplets below 32° F are said to be supercooled. At temperatures above 32° F, clouds are composed only of liquid droplets. At temperatures much below 15° F, they are usually composed mostly of ice crystals, while at intermediate temperatures they may be made up of supercooled droplets, ice crystals, or both.

Why dont ice crystals form more readily? First, the formation of ice crystals at temperatures higher than - 40° F requires sublimation nuclei. As was mentioned above, these usually are scarce in the atmosphere, especially at higher elevations. Also, many types of nuclei are effective only at temperatures considerably below freezing. However, another reason why vapor condenses into liquid droplets, rather than sublimes into ice crystals, is that condensation can begin at relative humidities well under 100 percent while sublimation requires at least saturation conditions and usually supersaturation.

Sublimation

Given the necessary conditions of below-freezing temperature, effective sublimation nuclei, and supersaturation discussed above, sublimation starts by direct transfer of water vapor to the solid phase on a sublimation nucleus. There is no haze phase as in the case of condensation. Once sublimation starts, ice crystals will grow freely under conditions of supersaturation. Since there are fewer sublimation than condensation nuclei available, the ice crystals that form grow to a greater size than water droplets and can fall from the base of the cloud.

Only very light snow, or rain if the crystals melt, can be produced by sublimation alone. Moderate or heavy precipitation requires one of the precipitation processes in addition to sublimation.

Precipitation

After condensation or sublimation processes have gone as far as they can, some additional process is necessary for droplets or crystals to grow to a size large enough to fall freely from the cloud and reach the ground as snow or rain. Cloud droplets, because of their small size and consequent slight pull of gravity, negligible rate of fall, and for all practical purposes are suspended in the air. Even drizzle droplets seem to float in the air. Raindrops range in size from about 1/50 inch to 1/5 inch in diameter. Drops larger than 1/5 inch tend to break up when they fall. It takes about 30 million cloud droplets of average size to make one raindrop about 1/8 inch in diameter.

There seem to be two processes which act together or separately to cause millions of cloud droplets to grow into a raindrop. One is the ice-crystal process and the other is the coalescence process.

The Ice-Crystal Process

We have seen that ice crystals and cloud droplets can coexist in clouds with subfreezing temperatures. For the ice-crystal process of precipitation to take place, clouds must be composed of both ice crystals and supercooled liquid cloud droplets.

Earlier we discussed vapor pressure and saturation vapor pressure at some length, but we considered only saturation vapor pressure with respect to liquid water. The saturation vapor pressure with respect to ice is somewhat less than that with respect to supercooled water at the same temperature.

If a cloud containing supercooled water droplets is saturated with respect to water, then it is supersaturated with respect to ice, and the relative humidity with respect to ice is greater than 100 percent. The force resulting from the difference between vapor pressure over water and over ice causes vapor molecules to be attracted to ice crystals, and the ice crystals will grow rapidly. As the ice crystals gather up vapor molecules in the cloud, the relative humidity with respect to water drops below 100 percent, and liquid cloud droplets begin to evaporate. Vapor molecules move to the ice crystals and crystallize there. Thus, the ice crystals grow at the expense of the water droplets and may attain a size large enough to fall out of the cloud as snowflakes. If the snowflakes reach warmer levels, they melt and become raindrops. This is the ice-crystal precipitation process.

Coalescence

Since rain also falls from clouds which are entirely above freezing, there must be a second precipitation process. This is a simple process in which cloud droplets collide and fuse together, or coalesce. Clouds which produce precipitation are composed of cloud droplets of varying sizes. Because of the different sizes, cloud droplets move about at different speeds. As they collide, some of them stick together to form larger drops. The larger cloud droplets grow at the expense of smaller ones, and actually become more effective in the collecting process as they become larger. As larger drops begin to fall, they tend to sweep out the smaller drops ahead of them.

The coalescence process takes place in clouds of both above freezing and below freezing temperatures. Snowflakes coalesce with other snowflakes as they fall to form the large clumps which we sometimes observe. They may also coalesce with supercooled water droplets to form snow pellets.

See also: Artificial Nucleation.

The knowledge that frequently there is a scarcity of sublimation nuclei and ice crystals in supercooled clouds has led to the discovery that precipitation can be initiated artificially. It has been found that silver-iodide crystals, which have a structure similar to ice crystals, can be effective sublimation nuclei in supercooled clouds at temperatures below about 20° F. Silver-iodide crystals can be released in the cloud by aircraft or rockets, or carried to the cloud by convection from ground generators.

Ice crystals can be created in a supercooled cloud by dropping pellets of dry ice, solid carbon dioxide, into the cloud from above. The dry ice, which has a melting temperature of -108° F, cools droplets along its path to temperatures lower than - 40° F so that they can freeze into ice crystals without the presence of sublimation nuclei. Once crystals are produced, they act as nucleating particles themselves and affect other parts of the cloud.

Once formed in a supercooled water cloud, ice crystals may grow by the ice-crystal process and coalescence, processes until they are large enough to precipitate. Studies have provided evidence that the artificial nucleation of supercooled clouds can, under the proper conditions, increase local precipitation significantly.

Precipitation products can be divided into three basic classes depending on their physical characteristics when they strike the earth: liquid, freezing, and frozen.

Liquid Precipitation

Rain and drizzle are the two kinds of liquid precipitation. The difference is mainly one of size and quantity of droplets. Drizzle droplets range in size from about 1/500 to 1/50 inch. Drizzle is formed in, and falls from, stratus clouds, and is frequently accompanied by fog and low visibility. Raindrops range in size from about 1/100 to 1/4 inch. They are much more sparse than drizzle droplets. Rain may come from liquid droplets formed by the coalescence process in warm clouds, or from melted snowflakes originally formed in cold clouds by both the ice crystal and coalescence processes. The snowflakes melt when they reach air with above-freezing temperatures. Rainfall intensity may vary from a few drops per hour to several inches in a matter of minutes. Heavier rainfall usually consists of larger drops.

Freezing Precipitation

Freezing rain and freezing drizzle are formed and fall as liquid drops that freeze on striking the ground. The drops may be above freezing, but usually they are supercooled and freeze upon striking the ground or other cold objects. This occurs usually with warm-front rain formed in the warm air above the frontal surface, and then supercooled as it falls through the cold air beneath the front. The temperature at the ground must be lower than 32° F.

Frozen Precipitation

Frozen precipitation consists of snow, snow pellets, sleet and hail.

• Snow consists of crystals of ice formed in pure ice clouds or in mixed clouds. The larger snowflakes are built up by the coalescence process. Air beneath the cloud must be near or below freezing, or the snow will melt before reaching the ground. The heaviest snowfalls occur when the temperature of the cloud portion from which the snow is falling is not much below freezing.

• Snow pellets are white opaque grains of ice, usually round. They form when ice crystals coalesce with supercooled droplets, and usually occur in showers before or with snow. They range in size from 1/16 to 1/4 inch.

• Sleet consists of transparent hard pellets of ice, about the size of raindrops, that bounce on striking the ground. They are formed by freezing of raindrops or by refreezing of partly melted snowflakes falling through a below-freezing layer of air. Sleet occurs most commonly with warm fronts.

• Hail consists of balls of ice ranging in size from 1/5 inch to several inches in diameter. They have layerlike structures indicating that they have grown by successive steps. Hailstones apparently begin their growth when supercooled water droplets impinge on ice pellets. The liquid water freezes on the ice pellet to form a layer of ice. This process is repeated until the hailstone falls out of the cloud. The repetition may be due to the hailstone being caught in strong updrafts and carried upward into the region of supercooled droplets. It is also possible for the process to begin at very high altitudes, in which case the hailstone grows as it falls through successive concentrations of supercooled water. Hail is associated with thunderstorms and very unstable air.

Dew and Frost

There are two other forms in which moisture from the atmosphere is deposited on the ground. These are dew and frost. Dew and frost do not fall, but instead are deposited when water vapor condenses or sublimes on the ground or on objects near the ground. Dew forms when air next to the ground or to cold objects is chilled to the dew point of the air, but remains above freezing. A common example is the deposit of water that forms on a glass of ice water. Frost forms by sublimation when the air is chilled to its dew point and the dew point is below freezing. Dew and frost forming on forest fuels at night can add considerably to the fuel moisture.

Precipitation is measured on the basis of the vertical depth of the water or melted snow. Snow, sleet, hail, and other solid forms are also measured on the basis of the depth of the unmelted form.

The standard rain gauge is an 8-inch cylindrical container with an 8-inch funnel at the top and a measuring tube inside. The cross-sectional area of the measuring tube is exactly one-tenth that of the funnel top. Thus, if 0.01 of an inch of precipitation falls, it is 0.1 inch deep in the measuring tube. The stick used to measure the precipitation is graduated in inches, tenths, and hundredths, so that 0.01 inch of rain is indicated for each 0.1 inch of stick length. When snow is measured, the funnel and measuring tube are removed, and only the outside cylindrical container is used. Snow caught in the gage is melted and measured in the measuring tube to obtain the liquid equivalent of the snow.

Several types of recording gages that make continuous records of the precipitation are also in use. The tipping bucket gage can be used only for rain. For each 0.01 inch of rain, an electrical impulse is recorded. Another type is the weighing-type gage, which can be used for either snow or rain. This device simply weighs the snow or rain that is collected. The weight is recorded continuously in inches of water on a chart attached to a revolving drum.

The rain gage should be exposed in the open away from large buildings or trees. Low bushes, fences, and walls are not objectionable, provided that the gage is placed at a distance of at least twice the height of the object. The top of the gage should be level.

In order to recognize and identify clouds it is necessary to classify and name them. Clouds are identified by their development, content and appearance, and their altitude. They are classified into many types and subtypes, but we need be concerned only with the more basic types. We will consider four families of clouds distinguished by their height of occurrence:

1. high clouds,
2. middle clouds,
3. low clouds, and
4. clouds with vertical development.

Within the first three families are two main subdivisions:

1. Clouds formed by localized vertical currents, which carry moist air upward beyond the condensation level. These are known as cumuliform clouds and have a billowy or heaped-up appearance.
2. Clouds formed by the lifting of entire layers of air, without strong, local vertical currents, until condensation is reached. These clouds are spread out in layers or sheets and are called stratiform.

In addition, the word nimbus is used as a prefix or suffix to indicate clouds producing precipitation--resulting in such names as nimbostratus or cumulonimbus. The word fractus is used to identify clouds broken into fragments by strong winds--such as stratus fractus and cumulus fractus.

Air stability has an important effect on the type of cloud formation. A stable layer which remains stable through forced lifting will develop stratiform clouds. Cumuliform clouds develop in air that is initially unstable or becomes unstable when it is lifted. A layer of conditionally unstable air which is forced to ascend may first develop stratiform clouds and then develop cumuliform clouds as the layer becomes unstable. The cumuliform clouds project upward from a stratiform cloud layer. Thus, the type of cloud formation can be used as an indicator of the stability of the atmospheric layer in which the clouds are formed.

High Clouds

High clouds have bases ranging from 16,500 to 45,000 feet. Cirrus, cirrocumulus, and cirrostratus clouds are included in this family. They are usually composed entirely of ice crystals, and this is their most distinguishing characteristic.

• Cirrus clouds are isolated wisps of thin, feathery cloud up near the top of the troposphere. They are sometimes called "mares tails" and may have trailing streams of larger ice crystals beneath them.
• Cirrocumulus clouds consist of patches of small, white cloud elements. They may form definite patterns at times, showing small but firm waves or ripples. They are sometimes referred to as "mackerel sky." True cirrocumulus are rare and are associated with other forms of typical cirrus at the same level, often changing into other forms of cirrus in a short time.
• Cirrostratus clouds are thin, whitish veils, sometimes covering the entire sky. Halos around the sun or moon, caused by their ice-crystal composition, frequently identify this cloud type.

Cirrus-type clouds may be produced in several ways. Often they are the forerunner of warm-front activity and give advance warning. Sometimes they are associated with the jet stream and usually are found on the south side of the jet. They may also be produced from the anvil tops of thunderstorms. The value of cirrus clouds in fire weather is their advance warning of warm-front activity and their use in indicating high-altitude moisture and wind direction and speed.

Middle Clouds

Middle clouds have bases ranging from 6,500 feet up to 20,000 feet. Altocumulus and altostratus clouds fall into this group. Middle are most generally formed by either frontal or orographic lifting, but may be formed in other ways. Often they are associated with lifting by convergence in upper-air troughs and sometimes develop with thunderstorms.

• Altocumulus are white or gray patches, with each individual component having a rounded appearance. Altocumulus may appear as irregular cloudlets or in definite patterns such as bands or rows parallel or at right angles to the wind. Usually, the stronger the wind, the more distinct the pattern. Altocumulus are usually composed of water droplets and often are supercooled.
• Altostratus appears as a gray or bluish layer or veil with a sort of fibrous texture. It may be made up of supercooled water droplets or a mixture of water droplets and ice crystals. It tends to cover the entire sky, and the sun will shine through dimly as through a frosted glass. As altostratus becomes thicker and lower, the sun becomes obscured. If it becomes dense and low enough, it becomes nimbostratus and takes on a wet and rainy appearance due to widespread precipitation. If the precipitation evaporates before reaching the ground, it is called virga.

Three special, types of middle clouds are of considerable importance in identifying weather conditions.

• The lens-shaped lenticular cloud appears over the ridge and to the lee of mountain ranges. Lenticular clouds indicate waves in the air flow caused by strong winds blowing across the range. The clouds form in the rising current on the upwind side of the wave crest and dissipate in the downward flow on the other side.
• The castellanus type of cloud consists of cumuliform masses in the form of turrets, usually arranged in lines. They indicate marked instability at high levels, and their occurrence in the forenoon is a warning of possible thunderstorm activity in the afternoon.
• Clouds with rounded lower surfaces in the form of pouches or udders are called mammata. They are most common on the underside of cumulonimbus anvils. The pendulous blobs of cloud are sinking into the clear air immediately below because they contain droplets and ice crystals. As the droplets and crystals evaporate, they chill the air in the pendants. By this chilling, they keep the air denser than the surrounding clear air. The process ends when all cloud particles have evaporated. Pilots have reported that the downdrafts within mammata are quite weak. However, any adjacent thunderstorms will have zones of marked turbulence.

Low Clouds

The bases of low clouds range from the surface to 6,500 feet. Low clouds include stratus, stratocumulus, and nimbostratus.

• Nimbostratus clouds form a gray, often dark, layer usually accompanied by continuously falling rain or snow. The precipitation usually reaches the ground, but occasionally only virga appears. Nimbostratus usually develops from thickening and lowering altostratus. Stratus fractus or scud clouds often appear beneath the nimbostratus layer.
• Stratus and stratocumulus are very common and widespread. They usually occur beneath an inversion and are fairly thin, ranging from a few hundred to a few thousand feet thick. Stratus forms a low, uniform sheet, dull gray in appearance. It is composed of water droplets and does not produce rain, although it may produce drizzle. Fog is simply a stratus cloud lying on the surface. When a fog layer lifts, as it frequently does during the forenoon, it becomes a stratus layer. In some localities, particularly the west coast, stratus is referred to as high fog.

Fog is important in fire weather because of its effect on the moisture content of forest fuels. While fog is forming or persisting, conditions are favorable for fuels to absorb moisture. Fog occurs during calm or light-wind conditions in a stable atmosphere and is formed in several ways. Radiation fog is formed when moist air cools to its dew point at night over a strongly radiating surface. Some vertical mixing is necessary to produce a layer of fog of significant thickness. Advection fog forms when warm, moist air passes over a cool surface and its temperature is reduced to the dew point. Many fogs are a combination of these two types. Upslope fog forms when moist, stable air is forced to rise along a sloping land surface. This type occurs especially along the western edge of the Great Plains when mT air moves northwestward from the Gulf of Mexico. Fog may also occur in connection with fronts, particularly in advance of warm fronts where evaporating rain falling through a layer of cold air near the surface saturates the cold air.

The distinction between stratus and stratocumulus is not particularly important. Stratocumulus shows individual rolls or rounded masses, usually soft and gray. It forms when the air is somewhat unstable, whereas stratus forms in stable air. Like stratus, it is composed of small water droplets and may produce light drizzle.

Clouds with Vertical Development

Clouds with vertical development include cumulus and cumulonimbus. These are irregularly shaped masses with domes or turrets and have a cauliflower appearance. They usually appear in groups, and individual cloud bases are at about the same altitude. The height of the bases, which is the condensation level, depends upon the air temperature and the amount of moisture in the atmosphere. Cumulus clouds are formed near the top of rising convection columns, and their bases may range in height from a few thousand feet to 15,000 feet or more. Their presence is of special interest in fire weather as an alert to possible convection in the surface layer. They are a common type during the fire season, particularly in mountainous regions.

The most common type of cumulus is a small, puffy type occurring during fair weather, called cumulus humili or fair weather cumulus. They appear after local surface heating becomes sufficiently intense to support convection, and dissipate in the late afternoon as surface heating decreases and convection ceases. These clouds have relatively flat bases, rounded or cone-shaped tops, and are usually isolated or in small groups. Their vertical growth is usually restricted by a temperature inversion, which makes the tops fairly uniform. Occasionally a single cloud element will develop vertically to some height. True fair-weather cumulus clouds, however, remain flat, but their presence indicates local updrafts that may influence fire behavior.

The danger from cumulus clouds is more acute, however, if the air is sufficiently moist and unstable to support their growth into towering cumulus. Virga or rain sometimes falls from the base of large cumulus.

The final stage of cumulus development is the cumulonimbus or thunderhead, which is characterized by a flat anvil-like formation at the top. The stretched-out shape of the anvil indicates the direction of air motion at that level. The anvil top is composed of sheets or veils of ice crystals of fibrous appearance, which are sometimes blown off to form cirrus-type clouds. Dissipating anvils give the appearance of dense cirrus and are sometimes referred to as false cirrus.

The greater the vertical development of cumulonimbus, the more severe the thunderstorm. Tops of cumulonimbus may extend to altitudes of 60,000 feet or higher and often reach the tropopause. Rain or snow showers usually accompany cumulonimbus clouds, and thunder, lightning, and hail are common.

Cumulus and cumulonimbus clouds not associated with frontal or orographic lifting indicate strong surface heating and atmospheric instability from the surface up through the level of the cloud tops. Surface winds are likely to be gusty and increase in speed as the cumulus forms. Other convection phenomena such as dust devils, whirlwinds, and considerable turbulence may be present. In addition to lightning, strong cold downdrafts present a threat from well-developed thunderheads (see: Thunderstorms).

Cumulus cloud caps often form atop the convection columns over large forest fires. Their moisture source may be almost entirely water vapor from the combustion process, or it may be water vapor entrained with air through which the column rises. Such clouds occasionally produce showers, but this is quite rare.

A thunderstorm is a violent local storm produced by a cumulonimbus cloud and accompanied by thunder and lightning. It represents extreme convective activity in the atmosphere, with both updrafts and downdrafts reaching high speeds. Thunderstorms are usually classified as frontal or air-mass thunderstorms. The thunderstorm depends upon the release of latent heat, by the condensation of water vapor, for most of its energy. Tremendous amounts of this energy are in a single well-developed thunderstorm. Part of the heat energy is converted to kinetic energy of motion to cause the violent winds, which usually accompany thunderstorms. Several conditions are necessary for thunderstorm development: a conditionally unstable atmosphere, sufficient moisture, and some lifting or triggering mechanism. Once initiated, thunderstorm cells go through a life cycle consisting of cumulus, mature, and dissipating stages. The most active stage is the mature stage when lightning discharges, the thunderstorm downdraft, and precipitation are all at their maximum. In extreme conditions, tornadoes may develop. Thunderstorms in the mature stage are important in fire control because they start fires by lightning, blow them out of control with the downdraft and outflow, or put them out with rain.

• Conditions Necessary for Thunderstorm Development
• Life Cycle of a Thunderstorm Cell
• Types of Thunderstorms
• Lightning
• Tornadoes

Thunderstorms have their origins in cumulus clouds. But only a few cumulus clouds develop into thunderstorms. Certain atmospheric conditions are necessary for this development to take place. These are:

1. Conditionally unstable air,
2. some triggering mechanism to release the instability, and
3. sufficient moisture in the air.

These factors may be present in varying degrees so that in one situation on a sultry afternoon only fair-weather cumulus will form, while in another situation numerous thunderstorms will develop. In the first situation, the instability in the lower atmosphere may be offset by stability aloft, which prevents strong convectional activity essential to the development of cumulonimbus clouds.

For thunderstorm formation, the air must be conditionally unstable through a deep layer. Convection must develop well beyond the freezing level for an electrical potential to be produced which will cause a lightning discharge. The conditional instability is released when the air is lifted to the level of free convection. Beyond this level, the lifted air is buoyant and rises freely and moist-adiabatically until it has cooled to the temperature of the surrounding air.

The triggering mechanism necessary to release the instability is usually some form of lifting. This lifting may be orographic or frontal, or may be produced by low-level converging flow or by heating from below. Any of these lifting processes may bring warm air from near the surface up to the level of free convection, above which it will rise freely.

Another triggering mechanism is the further steepening of the temperature lapse rate through advection of cold or warm air. Cold air moving in at high levels will steepen the lapse rate and make the atmosphere more unstable. Warm air moving in at low levels will have, the same steepening effect.

Clouds will not form in air containing little moisture even though other factors present may be favorable for thunderstorm development. For cumulus clouds to develop, air must be lifted to the condensation level, and for significant cloud growth it must be further lifted to the level of free convection. The greater the air moisture, the lower the condensation level and the easier it is for the level of free convection to be reached. Above the condensation level, the heat released in the condensation process tends to make the rising, air more buoyant. For this reason, the air need be only conditionally unstable rather than absolutely unstable for thunderstorms to develop when other factors are favorable.

The building upward of cumulus clouds into cumulonimbus may be prevented by layers of air at intermediate levels which are initially very stable or dry. Thunderstorms are unlikely to develop under these conditions even though all other factors favor development.

The thunderstorms that we see are composed of one or more individual convection cells. A storm composed of a cluster of cells will contain cells in various stages of development and decay. Each cell goes through a definite life cycle which may last from 20 minutes to 11/2 hours, although a cluster of cells, with new cells forming and old ones dissipating, may last for 6 hours or more.

Individual thunderstorm cells have many variations in growth and behavior, but typically go through three stages of development and decay. These are the cumulus, mature, and dissipating stages.

Cumulus Stage

The cumulus stage starts with a rising column of moist air to and above the condensation level. The lifting process is most commonly that of cellular convection characterized by strong updraft. This may originate near the surface or at some higher level. The growing cumulus cloud is visible evidence of this convective activity, which is continuous from well below the cloud base up to the visible cloud top. The primary energy responsible for initiating the convective circulation is derived from converging air below. As the updraft pushes skyward, some of the cooler and generally drier surrounding air is entrained into it. Often one of the visible features of this entrainment is the evaporation and disappearance of external cloud features.

The updraft speed varies in strength from point-to-point and minute-to-minute. It increases from the edges to the center of the cell, and increases also with altitude and with time through this stage. The updraft is strongest near the top of the cell, increasing in strength toward the end of the cumulus stage. Cellular convection implies downward motion as well as updraft. In the cumulus stage, this takes the form of slow settling of the surrounding air over a much larger area than that occupied by the stronger updraft. During this stage, the cumulus cloud grows into a cumulonimbus.

Cloud droplets are at first very small, but they grow to raindrop size during the cumulus stage. They are carried upward by the updraft beyond the freezing level where they remain liquid at subfreezing temperatures. At higher levels, liquid drops are mixed with ice crystals, and at the highest levels, only ice crystals or ice particles are found. During this stage, the raindrops and ice crystals do not fall, but instead are suspended or carried upward by the updraft (see also: Condensation, Sublimation, and Precipitation). Air temperature within the rapidly growing cell in this stage is higher than the temperature of the air surrounding the cell.

Surface weather during the cumulus stage is affected very little. Surface pressure falls slightly. Shade provided by the cloud during the daytime allows the ground to cool, and fuel temperatures approach that of the surface air. Except for cells, which develop above a frontal surface, the surface wind field shows a gentle inturning of winds forming the area of convergence under the updraft. The updraft at the center feeds into the growing cloud above. If a cloud with its updraft passes over a going fire, the convection from the fire may join with the updraft and they may reinforce each other. This joining may strengthen the inflow at the surface and cause the fire to become active.

Mature Stage

The start of rain from the base of the cloud marks the beginning of the mature stage. Except under arid conditions or with high-level thunderstorms, this rain reaches the ground. Raindrops and ice particles have grown to such an extent that they can no longer be supported by the updraft. This occurs roughly 10 to 15 minutes after the cell has built upward beyond the freezing level. The convection cell reaches its maximum height in the mature stage, usually rising to 25,000 or 35,000 feet and occasionally breaking through the tropopause and reaching to 50,000 or 60,000 feet or higher. The visible cloud top flattens and spreads laterally into the familiar "anvil" top. A marked change in the circulation within the cell takes place.

As raindrops and ice particles fall, they drag air with them and begin changing part of the circulation from updraft to downdraft. The mature stage is characterized by a downdraft developing in part of the cell while the updraft continues in the remainder. The air being dragged downward by the falling rain becomes cooler and heavier than the surrounding air, thus accelerating its downward fall. Melting of ice and evaporation of raindrops cool the descending air. The change from updraft to downdraft is progressive. The downdraft appears to start first near the freezing level and spreads both horizontally and vertically. The updraft continues in its decreasing portion of the cloud and often reaches its greatest strength early in the mature stage. The speed of the downdraft within the cell varies, but may reach 30 m.p.h. Usually it is not so strong as the updraft, which may exceed 50 m.p.h. The downdraft becomes most pronounced near the bottom of the cell cloud where the cold air appears to cascade downward.

Below the cloud, in the lower 5,000 feet or so above the ground, the downward rush of cool air decreases somewhat. The effect of a flat ground surface is to force the downdraft to pile up and spread out horizontally as a small, but intense, cold front. This horizontal outflow of air produces a strong and highly turbulent surge, frequently referred to as the "first gust." As this initial surge strikes an area, it causes a sharp change in wind direction and an increase in speed. This wind discontinuity is most pronounced on the forward side of the thunderstorm. Here, the storms movement is added to the speed of the outflow. To the rear, the storms movement opposes the outflow and makes it much less pronounced.

Because the outflowing air is cold and heavy, the first gust is accompanied by a sudden temperature drop, sometimes as much as 25° F, and a sharp rise in surface pressure. The pressure remains high as long as the dome of cold air is over an area.

The mature stage is the most intense period of the thunderstorm. There is extreme turbulence in and below the cloud, with intense gusts superimposed on the updraft and downdraft. Lightning frequency is at its maximum. Heavy rain and strong gusty winds at ground level are typical of most thunderstorms, though precipitation at the ground may be absent in high-level thunderstorms (see: Types of Thunderstorms). The heaviest rain usually occurs under the center of the cell, shortly after rain first hits the ground, and gradually decreases with time.

Dissipating Stage

As the downdrafts continue to develop and spread vertically and horizontally, the updrafts continue to weaken. Finally, the entire thunderstorm cell becomes an area of downdrafts, and the cell enters the dissipating stage. As the updrafts end, the source of moisture and energy for continued cell growth and activity is cut off. The amount of falling liquid water and ice particles available to accelerate the descending air is diminished. The downdraft then weakens, and rainfall becomes lighter and eventually ceases. As long as downdrafts and rain continue, temperatures within the cell are lower than in the surrounding air. As the downdrafts cease, air in the cell is gradually mixed with, and becomes indistinguishable from, the surrounding air. Then, either complete dissipation occurs or only stratiform clouds at lower levels and the separated anvil top remain.

As the thunderstorm cell dissipates, the surface signs also disappear unless new cells develop. Wind, temperature, and pressure gradually return to the conditions outside the thunderstorm area.

New Cell Development

Although each thunderstorm cell goes through a life cycle, different cells within a cluster at any time may be in various stages of development. As old cells die out, new ones are formed. The downdraft and outflowing cold air appear to be an important factor in the development of new cells. The preferred place for new cell development is the area between two cells where their outflowing cold air collides and causes upward motion in the overlying warm air. The forward edge of the cold dome may also act as a small cold front and cause lifting of warm air and the development of new cells. Local topographic features may also influence the initiation of new cells. A cell may form over a mountain peak and drift off downwind as another cell develops over the peak.

The interaction of cells in a cluster can cause false impressions of the behavior of thunderstorms. Thunderstorm cells usually move in the direction of the airflow in the layer in which they develop but at a speed somewhat less than this airflow. Cell growth, decay and replacement of old cells, and the extension of the storm area by new cell formations may make the storm system appear to split, back into the wind, turn at right angles to the wind, or move faster than the general wind itself. The true movement is difficult to discern from the ground, particularly in mountain topography.

Thunderstorms are usually classified as frontal or air-mass thunderstorms.

Frontal Thunderstorms

Frontal thunderstorms are caused by warm, moist air being forced over a wedge of cold air. This lifting may occur with warm fronts, cold fronts, or occluded fronts.

• Warm-front thunderstorms are usually embedded in large stratiform cloud masses. They are likely to be the least severe of frontal thunderstorms because of the shallow slope of the warm-front surface. Surface wind conditions, in the cold air wedge beneath the warm front, may be unaffected by the thunderstorms above.
• Cold-front thunderstorms are generally more severe and occur in a more-or-less continuous line. Their bases are normally lower than those of other frontal thunderstorms.
• Thunderstorms occurring along a squall line are similar to those along a cold front, but may be even more severe. Heavy hail, destructive winds, and tornadoes are usually associated with squall-line thunderstorms.
• Thunderstorms are often associated with a warm-front type occlusion. In this case, they occur along the upper cold front and are set off by the lifting of the warm, moist air. They are usually more severe than warm-front thunderstorms and less severe than the cold-front type.

Air-mass Thunderstorms

Air-mass thunderstorms are unaffected by frontal activity. They are usually scattered or isolated. Air-mass thunderstorms may be further classified as convective or orographic, although these lifting processes often act together.

• Convective thunderstorms formed by convergence may occur day or night, but they tend to be most active in the afternoon. Those produced by instability resulting from advection of low-level warm air or high-level cold air may also occur day or night. The nocturnal, or nighttime, thunderstorm, which is common in the Midwest during spring and summer, is usually due to low-level warm-air advection and convergence. These storms are among the most severe found anywhere.
• Orographic thunderstorms develop when moist, unstable air is forced up mountain slopes. They tend to be more frequent during the afternoon and early evening because heating from below aids in the lifting process. Storm activity is usually scattered along the individual peaks of mountain ranges, but occasionally there will be a long unbroken line of thunderstorms.

High-level or Dry Thunderstorms

One type of air-mass thunderstorm, the high-level or dry thunderstorm, deserves special consideration because of its importance in starting wildfires. The lifting process may be orographic, convergence, cold-air advection aloft, or a combination of these, often aided by surface heating over mountain ranges. High-level thunderstorms occur most frequently in the mountainous West during the summer months.

Their distinctive feature is that their cloud bases are so high, often above 15,000 feet, that precipitation is totally or mostly evaporated before it reaches the ground. As a result, lightning strikes reaching the ground frequently start fires in the dry fuels. The downdraft and outflow usually reach the ground even though the precipitation does not. The cold, heavy air is generally guided by the topography into downslope and downcanyon patterns, but crossslope flow may also occur.

There are two principal weather patterns which produce high-level storms. One is the inflow of moist air, usually from over the Gulf of Mexico but occasionally from over the eastern subtropical Pacific, at levels of 10,000 to 18,000 feet. Thunderstorms are set off by lifting over mountains, and by heating and upslope thermal winds at higher levels in the mountains, as the moist air spreads northward from New Mexico, Arizona, and southern California. These storms usually develop in the afternoon and may extend into the evening hours.

The second important weather pattern in high-level storms is the cold Low aloft. With this pattern a closed low-pressure system aloft becomes cut off from the main belt of westerlies. The cold air within this closed Low produces instability and causes convective currents to develop. If sufficient moisture is present, thunderstorms will form. They can develop at any time of the day or night, but are most active in the afternoon when they are assisted by daytime heating. The movement of a closed upper Low is erratic and very difficult to predict. The Low may move in virtually any direction, may deepen or fill, or may be picked up by a trough moving eastward at a higher latitude.

The Far West is a favorite place for closed Lows to develop. They may meander around for several days or a week before finally dissipating or moving on.

Lightning occurs in a thunderstorm when an electrical potential builds up that is strong enough to exceed the resistance of the atmosphere to a flow of electrons between the centers of opposite charge. Most cloud-to-ground discharges originate in the cloud and progress to the ground. They take place in two stages. First, a leader stroke works its way downward to the ground in a series of probing steps. Then a number of return strokes flash upward to the cloud so rapidly that they appear as a flickering discharge. The average number of return strokes in a lightning flash is four. Lightning discharges taking place within a cloud usually do not show return strokes.

The processes that generate the electrical potential are not fully understood, and a number of theories have been advanced. Regardless of the method or methods by which electrical potentials are generated, measurements with specialized electronic equipment have established where in the thunderstorm opposite charges tend to accumulate and how charges vary during storm development.

In fair weather, the atmosphere has a positive electrical charge with respect to the earth. This fair weather potential gradient has an average value of about 30 volts per foot. When a cumulus cloud grows into a cumulonimbus, the electric fields in and near the cloud are altered and intensified. The upper portion of the cloud becomes positively charged and the lower portion negatively charged, although other smaller positive and negative charges develop. The negative charge near the cloud base induces a positive charge on the ground--a reversal of the fair-weather pattern.

Cloud-to-ground lightning is usually a discharge between the negative lower portion of the cloud and the induced positive charge on the ground and accounts for about one-third of all discharges. Most lightning discharges, however, are within a cloud or cloud-to-cloud. Many of the within-cloud discharges take place between the negative charge in the lower portion of the cloud and a positive charge center carried downward from the upper portion of the cloud by the falling rain in the precipitation core. This positive charge center disappears when the heavy rain stops.

Lightning sometimes occurs in the cumulus stage, but reaches its greatest frequency at the time the cell reaches maturity and its greatest height. The start of rain beneath the cloud base at the beginning of the mature stage marks the onset of the greatest lightning danger. The most extensive horizontal flashes occur at altitudes extending from the freezing level upward to where the temperature is about 15° F. Although lightning may occur throughout a thunderstorm cell, the strongest flashes to the earth usually originate in the lower portion of the cell. Many cloud-to-ground lightning strikes reach out laterally for considerable distances from the cloud base. Once lightning has started, it may continue well into the dissipating stage of the cell. Apparently, less cloud height is needed to maintain continuing discharges than to initiate the first. But as the height of the cell decreases after reaching maturity, the frequency of lightning flashes decreases. However, individual flashes may remain strong.

The noise of thunder is due to compression waves resulting from the sudden heating and expansion of the air along the path of the lightning discharge. These compression waves are reflected from inversion layers, mountainsides, and the ground surface so that a rumbling sound is heard, instead of a sharp explosive clap, except when the discharge is very near. Since light travels so very much faster than sound, it is possible to estimate the distance of a lightning flash using the elapsed time between seeing the flash and hearing the thunder. The distance to a flash is about 1 mile for each 5 seconds of elapsed time.

Weather radar, in which portions of transmitted radio signals are reflected back from precipitation areas in clouds and displayed as radar echoes on an indicator, is helpful in locating, tracking, and revealing the intensity of thunderstorms and their associated lightning.

Tornadoes are violent whirling storms which may occur with severe thunderstorms. They take the form of a funnel or tube building downward from a cumulonimbus cloud. These violently rotating columns of air range in size from a hundred feet to a half mile in diameter. Technically, they are not tornadoes unless they touch the ground, but are referred to as "funnel clouds." When they do reach the ground, they are the most destructive of all atmospheric phenomena on the local scale. They travel with a speed of 25 to 50 m.p.h., usually from southwest to northeast, and often skip along. The length of the path of a single tornado is usually just a few miles, but some tornadoes have remained active for more than a hundred miles-striking the ground for a few miles, skipping an area, then striking the ground again, and so on.

The great destructiveness of tornadoes is caused by the very strong wind and extremely low pressure. Winds in the rapidly spinning vortex have never been measured, but from the destruction it is estimated that winds may exceed 500 m.p.h. The low pressure causes houses and structures to virtually explode when a tornado passes over them. There is a sudden decrease in pressure around the house, while on the inside the pressure changes little. The resulting difference in pressure between the outside and the inside is sufficient to blow the house apart.

Tornadoes have been reported in all of the 48 contiguous States and Southern Canada, but they are rare west of the Rocky Mountains. Maximum occurrence is in the central Midwest, and there is a secondary maximum in the Southeast. In Southern United States tornadoes may occur in any month of the year, but farther north the maximum occurrence is in late spring and early summer. They generally occur with prefrontal squall lines, but they may develop with other violent thunderstorms, including those in hurricanes. Tornadoes usually occur in the late afternoon or evening. Their main effect on the wildland fire problem is the resulting blowndown timber in forested areas that often creates high fire hazard.

The fire weather occurring on a particular day is a dominant factor in the fire potential on that day. Fire climate, which may be thought of as the synthesis of daily fire weather over a long period of time, is a dominant factor in fire-control planning. In a broad sense, climate is the major factor in determining the amount and kind of vegetation growing in an area, and this vegetation makes up the fuels available for wildland fires.

The areas of North America in which wildland fires are a problem have a wide variety of fire climates. Latitude alone accounts for major changes from south to north. The shape of the continent, its topography, its location with respect to adjacent oceans, and the hemispheric air circulation patterns also contribute to the diversity of climatic types by affecting general temperature and precipitation patterns. In general, the fire season in the western and northern regions of the continent occurs in the summertime. But the fire season becomes longer as one goes from north to south, becoming nearly a year-round season in the Southwest and southern California. In the East, the fire season peaks in the spring and fall. Some fires occur during the summer months, and in the Southern States they can occur in winter also.

Because of the nature of the effects of various weather elements on fire behavior, simple averages of the weather elements are of little control value. Also, fire climate cannot be described by considering the weather elements individually. Fire potential responds to the combined effects of all of the fire-weather elements. For example, it makes considerable difference in fire climate whether or not precipitation is concentrated in the warm season or the cold season of the year. If it is concentrated in the cold season, and the warm season is dry, the fire potential during the warm season may be extreme. Where the reverse is true, the warm season may have little fire potential, while the most critical periods may be in spring and fall. Strong winds are very important in fire behavior, providing they occur in dry weather. A region may often have strong winds, but if they occur with precipitation, they are of much less importance to the fire climate.

Climatic differences create important variations in the nature of fire problems among localities and among regions-- seasonally and between one year and another. Knowledge of the similarities, differences, and interrelationships between regional weather patterns becomes a useful daily fire-control management device. A weather pattern that is significant to fire behavior in one region may be unimportant in another. What is unusual in one region may be commonplace in another. On the other hand, many large-scale weather patterns ignore regional boundaries, and one originating in or penetrating a region may then be a forewarning of what is soon likely to happen in neighboring regions. Fire-danger rating is an integration of weather elements and other factors affecting fire potential. Daily fire-danger rating is dependent on current fire weather, while seasonal and average fire-danger ratings are dependent on the fire climate. In many systems, only the weather elements are considered, because they are the most variable.

• Topographical Features of North America
• Maritime Influences in North America
• Pressure and General Circulation Patterns
• Temperature Variations
• Precipitation Patterns
• Fire Climate Regions
• Southeastern States
• Central States
• North Atlantic
• Great Plains
• Great Lakes
• Northern Rocky Mountains
• Southern Rocky Mountains
• Great Basin
• Southwest
• North Pacific Coast
• South Pacific Coast
• Interior Alaska and the Yukon
• Central and Northwest Canada
• Sub-Arctic and Tundra
• Mexican Central Plateau

The extent of the North American Continent in both its north-south and east-west dimensions permits the full development of continental air masses over much of the land area. The continent is also surrounded by water and is invaded by various maritime air masses. How both types combine to influence the North American climate is largely determined by the surface configuration of the land mass.

It is particularly important that only about a quarter of North America is covered by significant mountain topography. Furthermore, all the mountains lie on the far western side of the continent except two mountain chains along the Atlantic and Gulf of Mexico seaboards. These two chains are the Appalachian Mountains in the United States and the Sierra Madre Oriental in Mexico. It is also important that, with the exception of the Brooks and associated ranges enclosing interior Alaska and adjoining Canada, all of the major mountain systems have a north-south orientation.

The entire west coast is rimmed by a series of coastal ranges extending, with only infrequent interruptions, from southern Lower California to southern Alaska. A narrow coastal plain separates the mountains from the sea over most of this coastline from Mexico to southern British Columbia. From there northward, the Coast Mountains more commonly rise abruptly from near the water. Glaciers are common along the Canadian and Alaskan coasts, increasing in number northward.

Two disconnected interior ranges in the Far West have additional influences on climate. The Sierra Madre Occidental in Mexico, east of the Gulf of California, is a secondary range largely shielded from direct Pacific influence by the mountains of Lower California. South of the tip of Lower California, it becomes the mainland western coast range of Mexico. The Sierra-Cascade Range, beginning in the north portion of southern California, parallels the Coast Range up to the Fraser River in southern British Columbia. This range bounds the east side of California's Central Valley and a succession of coastal valley systems through Oregon and Washington. It is somewhat higher than the Coast Range, including several peaks in excess of 14,000 feet in elevation.

The Rocky Mountain system forms the backbone of that portion of the continent lying in Canada and the United States. It is the continent's most massive mountain expanse and forms the Continental Divide, separating water that flows to the Pacific from that flowing to all other surrounding waters. The mountains extend from the Arctic Ocean west of the Mackenzie River to northern New Mexico. The Sierra Madre Occidental plays a similar role in Northern Mexico.

The vast intermountain region west of the Rocky Mountains and northern Sierra Madre is known as the Cordilleran Highlands. From a narrow beginning in northern British Columbia, it extends southward in a generally broadening belt to Northern Mexico, where it becomes the Mexican Plateau, and diminishes in width farther south. In the United States a large part is called the Great Basin. The region, as its name implies, is upland country. Because of both topographic and latitudinal differences, however, there are some sub-regional characteristics that are also important to the climatology of the region as a whole.

East of the Rocky Mountains, all of Canada and parts of the Northern United States were scraped and gouged by the prehistoric Polar lee Cap. This left a land of many lakes and low relief covered mostly by glacial till and numerous moraines. This glaciated region extends into, and connects with, the broad Mississippi Valley system and the adjoining Great Plains-which slope upward to the foot of the Rocky Mountains from Southern Canada to Texas.

East of the Rocky Mountains, therefore, the Appalachian Mountains represent the only topographic barrier on the continent that has a significant influence on general air circulation. It is particularly noteworthy that there is no such barrier between the Arctic and the Gulf of Mexico.

The interiors of Canada and Alaska are source regions for continental polar air and are protected from maritime influence by the western mountain chains. Upon leaving the source regions, this cP air can penetrate far to the south because of the absence of any major east-west mountain ranges across the continent. The southflowing cold air is channeled between the Rocky Mountain system and the less formidable Appalachian Mountains. It often reaches and sometimes crosses the Gulf of Mexico. The lack of mountain barriers also allows warm, moist air from the Gulf of Mexico to flow northward. This warm air constitutes a somewhat deeper layer than the continental air and is less influenced by the mountain systems.

Because of its generally high elevation, the interior of Northern Mexico is little affected by polar continental air. Maritime influence is also restricted. The Sierra Madre Occidental in the west limits the surface effects of Pacific maritime air to the coastal strip. The Sierra Madre Oriental limits the surface effects of Gulf air to the coastal plan. As Mexico's land mass narrows toward the south between the adjacent warm Pacific and Gulf waters, the climate becomes warm and humid.

1. The typical northeast trade winds blowing onshore from the Atlantic and Gulf between the Tropics and 30° N,

2. prevailing westerlies off the Pacific between 30° N and the Polar Front zone, and

3. polar easterlies north of the Polar Front zone. As seasonal heating and cooling change, these pressure and wind systems move somewhat north in summer, and south again in winter.

Temperatures vary with the intensity of solar radiation at the earth's surface, among other factors. Because of this, there is a close relationship between average temperatures and latitude. Another major influence on temperature patterns is the distribution of land and water surfaces. At any given latitude, mean temperatures are higher in summer and cooler in winter over land than over water. The annual range of temperature between winter and summer is greater in the interior of the continent than over the adjacent oceans. The blocking effect of the high western mountain ranges also influences the mean temperature pattern.

A third major influence on temperature is elevation because the temperature through the troposphere usually decreases with height. Thus, an area a few thousand feet above sea level may have average maximum temperatures comparable to a low elevation area many hundreds of miles farther north.

Mean winter temperatures are higher along the east and west coasts than they are in the interior, and higher along the west coast than the east coast. These differences are more marked at higher latitudes than at lower latitudes. In the general west-to-east airflow, the west coast is more strongly influenced by the adjacent ocean than the east coast. In addition, the west coast is sheltered from the cold continental air masses by high mountain ranges.

In January, almost all of the interior of Canada and the Northern United States have mean temperatures below freezing. The coldest temperatures are found in the region between Hudson Bay and northern Alaska. The Great Lakes have a slight moderating effect on the temperature pattern; this area shows slightly higher mean temperatures than points to the east or west.

In the summer, differences in temperatures between the northern and southern sections of the continent are much less than in winter. The effect of the lesser angle of the sun's rays in the northern latitudes is partially offset by the longer days there. The sharp temperature gradient across the Pacific coastline is largely due to the cool California Current off the coast and the intense daytime heating which is felt, not only in the American Southwest, but also to some extent up through British Columbia and into interior Alaska.

The highest temperatures in summer are found in the lowlands of the Southwest; the lowest temperatures are found in Northeastern Canada.

In general, autumn temperatures are higher than spring temperatures in North America. There are some exceptions; in Texas and the interior of British Columbia, temperatures are higher in April than in October.

Both annual precipitation and seasonal distribution of precipitation depend on: (1) The moisture content of the air and vertical motions associated with surface heating and cooling, (2) major pressure systems, and (3) frontal and orographic lifting. This lifting has its greatest effect when the prevailing moist wind currents blow across major mountain systems.

In North America, the greatest precipitation is on the Northern Pacific coastal plains and the western slopes of the mountains, due to the influx of moist air from the Pacific Ocean. Maximum fall is on the Pacific Northwest coast of the United States, with amounts decreasing both north and south of this region. The inland valleys receive less precipitation than the coastal plains and coastal mountains. Along the western slopes of the next major ranges, such as the Sierra-Cascades, further lifting of the moist air again causes an increase in the total precipitation. A third, and final, lifting of these westerlies occurs on the western slopes of the Rocky Mountains, which extract most of the remaining precipitable moisture.

In each of these cases of orographic lifting, there is a decrease in precipitation activity as the air flows across the crests. Previous precipitation has left the air less moist. The lifting force has ceased, and often there is subsidence on the leeward side, which further reduces the degree of saturation. Such a leeward area is said to lie in a rain shadow, a term derived from its similarity to the shadows cast by the western mountains as the sun goes down. This explains why the inland valleys receive less precipitation than the coastal plains and mountains. The Great Basin area in the United States lies in such a rain shadow, and ranges from semi-desert to desert.

East of the Rocky Mountains, air of Pacific origin has become relatively dry, and its importance as a source of precipitation is replaced by moist air from the Gulf of Mexico. The influence of Gulf air extends northward well into Canada. Annual precipitation increases to the east and south under the more frequent intrusions of moist air from the Gulf and the Atlantic. The greatest annual precipitation is along the Gulf coast and the southern end of the Appalachians.

In most areas of the continent, there is considerable variation in annual rainfall. Wet and dry years may occur irregularly in poorly defined patterns, or as wet and dry fluctuations of variable duration. Within any one climatic region, a characteristic variation can usually be identified. Common ones are: Normally moist but with occasional critically dry years; typically dry with only infrequent relief; or longer period fluctuations of alternating wet years and dry years. The seasonal distribution of precipitation varies widely over the continent and is often as important in fire weather as the total annual amount.

Considering geographic and climatic factors together, it is possible to delineate 15 broad climatic regions over the North American continent. Most of these differ in one or more aspects, giving each a distinctive character affecting the wildland fire problem. In considering the climatic characteristics of a particular region, we should remember that generalities must be made and that there are many local exceptions.

• Southeastern States
• Central States
• North Atlantic
• Great Plains
• Great Lakes
• Northern Rocky Mountains
• Southern Rocky Mountains
• Great Basin
• Southwest
• North Pacific Coast
• South Pacific Coast
• Interior Alaska and the Yukon
• Central and Northwest Canada
• Sub-Arctic and Tundra
• Mexican Central Plateau

The vegetation in the Southeastern United States consists mainly of pines along the coastal plains, hardwoods in bays and bottomlands along stream courses, and mixed conifers and hardwoods in the uplands. Flash fuels, flammable even very shortly after rain, predominate in this region. The topography along the Gulf and Atlantic is low and flat. Inland from the Atlantic Coast it merges with an intermediate Piedmont area. The southern Appalachians are included in this region, and the central portion includes the lower Mississippi Valley.

Summers are warm and generally humid, because the region is almost continuously under the influence of an mT air mass. Winters have fluctuating temperatures. When mT air moves over the region, high temperatures prevail. Following the passage of a cold front, cP air may bring very cold temperatures--well below freezing--throughout the Southern States.

Annual precipitation varies from 40 to 60 inches over most of the region, except for about 70 inches in the southern Appalachians and over 60 inches in the Mississippi Delta area, and falls mostly as rain. The influence of the moist mT air from the Gulf of Mexico causes abundant rainfall in all seasons, with slightly higher amounts in August and September due to the presence of hurricanes in some years. Spring and fall have less precipitation than summer or winter, with spring being wetter than fall. Winter precipitation is usually associated with frontal lifting or with Lows that develop over the Southern States or the Gulf of Mexico and move through the region. Summertime precipitation is mostly in the form of showers and thunderstorms. During the colder months, much fog and low stratus are formed by the cooling of mT air as it moves northward.

The fire season in the Southern States is mainly spring and fall, although fires may occur during any month.

The four synoptic types that bring high fire danger to the other regions east of the Rockies also bring high fire danger to the Southern States. The Hudson Bay High and Northwest Canadian High types affect this region less often than the regions to the north. The airflow pattern aloft must have considerable amplitude for Highs from Canada to reach the Southern States.

The Pacific High type causes more days of high fire danger than any other type. Pacific Highs may reach this region with either meridional or zonal flow aloft. Very often, the most critical fire weather occurs with the passage of a dry cold front. The air mass to the rear may be mP or cP. The strong, gusty, shifting winds with the cold front and dry unstable air to the rear set the stage for erratic fire behavior.

The Bermuda High type is second to the Pacific High in causing high fire danger in this region. This type is rather stagnant and persists over the region for long periods of time, mostly in spring, summer, and fall. The cutting off of Gulf moisture by the Bermuda High, when it extends westward across the Southern States to Texas, is the typical drought pattern for this region. Aloft, a long-wave ridge is located over the central part of the continent and the belt of westerlies is far to the north, near the Canadian border.

Subsidence and clear skies produce low humidities and high temperatures. These factors, plus the extended drought, set the stage for high fire danger. Peaks in fire danger occur as winds increase with short-wave trough passages and their associated surface cold fronts on the north side of the Bermuda High. Lightning accounts for only a minor number of fires.

The vegetation in this region consists of grass in the lowlands, brush at intermediate levels, and extensive coniferous stands in the higher mountains.

Temperatures along the immediate coast are moderated both winter and summer by the ocean influence. But only short distances inland, winter temperatures are somewhat lower and summer temperatures average considerably higher.

The annual precipitation is generally light, around 10 to 20 inches at lower elevations. Precipitation in the mountains ranges up to 60 inches or more locally. Summers are usually rainless, with persistent droughts common in southernmost sections. Widespread summer thunderstorms, with little precipitation reaching the ground, particularly in the mountains of the northern half, occasionally result in several hundred local fires within a 2- or 3-day period.

The fire season usually starts in June and lasts through September in the north, but in the south critical fire weather can occur year round.

Several synoptic weather types produce high fire danger. One is the cold-front passage followed by winds from the northeast quadrant -the same as was described above for the coastal region farther north. Another is similar to the east-wind type of the Pacific Northwest coast, except that the High is farther south in the Great Basin. This Great Basin High type produces the foehn-type Mono winds along the west slopes of the Sierras and Coast Ranges, and the Santa Ana winds of southern California. Peak Santa Ana occurrence is in November, and there is a secondary peak in March. A third high fire-danger type occurs when a ridge or closed High aloft persists over the western portion of the United States. At the surface, this pattern produces very high temperatures, low humidities, and air-mass instability.

This is a region of rain-forest types with heavy coniferous stands. Because of the maritime influence, coastal areas are comparatively warm throughout the winter. The lowest temperatures occur when a cP air mass crosses the coastal mountains and covers the Pacific coast, but this is a rare event. Summer temperatures are rather cool, again because of the Pacific Ocean influence. There is a high frequency of cloudy or foggy days throughout the year.

The rainfall in this region is mostly concentrated in the winter months; summer rainfall is usually very light.

Annual rainfall varies from 60 to 150 inches along the coast, averaging 60 to 80 inches along British Columbia and the south Alaska coastal plains, 80 to 100 inches along the Pacific Northwest coast, and as low as 20 to 30 inches in some northern California coastal sections. Many local areas along the coastal slopes have much greater totals, with some areas receiving over 150 inches; in the Olympic Mountains, annual precipitation ranges up to 240 inches. The valley systems to the east of the Coast Ranges receive 12 to 20 inches in British Columbia, 30 to 50 inches in Washington and Oregon, and 15 to 20 inches in northern California.

The combination of high rainfall and moderate temperatures results in a buildup of extremely heavy fuel volumes. The maritime influence, particularly along the immediate coast, usually holds the fire danger to moderate levels during most seasons. However, some summers are very dry and warm with high fire danger. During these periods, fires are characterized by high intensities, firewhirls, and long-distance spotting.

The fire season usually runs from June through September. Lightning fires increase in number and severity from the coast inland.

In northern California and in western Oregon and Washington, strong, dry north to east winds may produce extreme fire danger in late summer and early fall. Two synoptic weather types produce this critical fire weather. One is a cold-front passage followed by a bulge of the Pacific High extending inland over the coast. The attendant northeasterly winds blowing downslope produce a warming and drying foehn effect. The second type follows when higher pressure develops east of the Cascades at the time a trough lies along the coast. The resulting dry easterly winds will cause high fire danger west of the Cascades. Airflow from the northeast quadrant not only keeps the marine air offshore, but also results in adiabatic warming as the air flows from higher elevations down to sea level.

The vegetation in this regain is predominantly spruce and aspen, with some tundra and other lesser vegetation in the north. The Yukon Basin has a warm, short summer. Continental heating has produced summertime temperatures of 100° F., but temperatures as low as 29° F. also have occurred in July. Winters are extremely cold. The high coastal mountains generally prevent the invasion of mP (maritime polar) air masses at low levels. The Brooks and other ranges block the inflow of even colder cP (continental polar) air from the north.

Annual precipitation is only about 10 to 15 inches, the maximum occurring during the summer in convective showers and with weak fronts. Precipitation is highest in the southern portion, which includes the northern extension of the Cordilleran Highlands and their parallel chains of lesser mountains. Although precipitation is maximum in summer, it is so scant that wildland fuels dry out considerably during the long, clear, dry summer days. Dry thunderstorms are not infrequent.

The usual fire season starts in May after melting of the winter snows and lasts until September.

With the exception of the southern prairies, vegetation in this part of Canada consists predominantly of spruce, pine, poplar, and aspen forest with various mixtures of other species. In spite of the short growing season in the far northwest, comparatively good tree growth results from the long daylight hours. Much of the vegetation in the region reflects an extensive past fire history.

This region is glaciated with mostly low relief, except for the more broken topography of the mountain foothills along the western boundary. A common characteristic is very low winter temperatures. The region serves as both a source region and southward pathway for cold cP air masses. The large north-south and east-west geographical extent of the region results in significantly different summer temperature and moisture regimes from one part of the region to another.

The far northwest portion of the region has long, predominantly clear, sunny days contributing to rapid and extensive drying of forest fuels. Even though the summer season is short, drying is only occasionally and temporarily alleviated by summer showers. Proceeding southward and eastward, the summer days are not as long, although the season is longer. On clear days, maximum temperatures may be considerably higher here than in the northwest portion of the region, but cloudy days with shower activity are frequent.

Precipitation distribution is an important part of the regional climatology. The average annual amounts vary from 8 to 10 inches in the far northwest, to 20 inches in southern portions of the Prairie provinces, and up to 30 inches at the eastern extremity. Winter snows are generally light because the cold air holds little moisture, so it is usual for at least half of the total precipitation to come in the form of summer rains. These rains often are thunderstorms with accompanying lightning fires, and they occur with varying frequencies in virtually all parts of the region. The principal cyclone tracks during the summer run through the central part of the region.

The geographic extent of this region is so great that it is not practical to designate any particular fire season for the area as a whole. For example, locally there may be both a spring and fall fire season, a summer fire season, or any combination of these.

This region, extending from the Mackenzie Delta to the Atlantic, supports scattered patches of scrub spruce forest in the south merging with open tundra in the north. It is all low glaciated terrain. Annual precipitation is about 10 to 15 inches in the northwest and up to 20 to 25 inches in the east. More precipitation falls in the summer than in winter.

The fire season is principally during midsummer. Strong winds and low humidities are common. The average number of fires is small, with apparently half or more caused by lightning. There is considerable evidence of severe past fire history.

The vegetation in the plateau region of Mexico is largely brush and grass with ponderosa pine at higher elevations. The region is a high plateau and mountainous area, generally above 6,000 feet, lying between the two principal north-south mountain ranges. It differs from the Southwest mainly in that it is affected more directly by moist air from both the Gulf of Mexico and the Pacific, although this influence is restricted by mountain barriers. Temperatures are comparatively cool for the latitude because of the elevation. Characteristically, the summers are warm with frequent convective showers and generally high humidities. The winters are cool and dry.

The annual precipitation is low to moderate. The maximum occurs in the summer with frequent thunderstorms due to continental heating. In spite of greater precipitation, the fire season is mostly in the summer.

The vegetation in the Central States region is mostly hardwoods, and mixed pine and hardwoods, interspersed with agricultural lands. The topography is mostly flat to gently sloping. The principal exceptions are the Missouri and Arkansas Ozarks and the western portions of the Appalachians. Summer temperatures tend to be high in the southern portion of the region, but relative humidities are usually high also. The northern portion experiences brief periods of high temperatures and brief periods of moderate temperatures as mT air masses alternate with either mP or cP air masses. Winters can be extremely cold in the north.

Annual precipitation is moderate, generally 20 to 45 inches, with snow and rain in the winter, and showers and thunderstorms in the summer. Usually, there is sufficient rain with thunderstorm activity to minimize lightning fire occurrence. The maximum precipitation usually falls in early summer in the north, but there is a fair distribution throughout the year in the southern portion. There are occasional dry summers, but the green tree canopies and green lesser vegetation are usually sufficiently effective in the summer to keep fires from being aggressive.

As in the Great Lakes region, the principal fire season is in spring and fall when the hardwoods are not in leaf and the lesser vegetation is dead. In the southern portion of the region the spring season is somewhat earlier and the fall season somewhat later than in the northern portion.

The synoptic weather patterns producing high fire danger in the Central States are similar to those affecting the Great Lakes region, except that the Bermuda High type influences the southern portion of the Central States region more frequently. Nevertheless, the Bermuda High is the least important of the types, both from the standpoint of frequency and from the fact that it occurs mainly during the summer months when vegetation is green. The Pacific High, Hudson Bay High, and Northwestern Canadian High types, in that order, cause nearly all of the high fire danger in spring and fall. These types have been described above for adjoining fire climate regions.

The forests in the North Atlantic region vary from extensive spruce stands in the north to predominantly hardwoods in the southern portions. The region is bounded on the west by the crest of the Appalachians and on the east by the sea. The coastal plain is wider than that facing the Pacific and increases in width from north to south. The immediate coast is influenced by the Atlantic Ocean and often is cool and foggy. But because the general movement of weather systems is from west to east, the maritime influence usually does not extend far inland. For this reason, temperatures can be quite low in winter and quite high in summer. On occasion, mP air from the Atlantic moves sufficiently southwestward to influence this region.

The annual precipitation is moderate to heavy, with totals of 40 to 50 inches, and is fairly well distributed throughout the year. There is a slight maximum during the summer and a slight minimum during the spring. Storms moving into the region from the west do not produce as much precipitation on the east side of the mountains as storms which move northeastward along the coast. In the first case, the descending flow on the east side of the mountains diminishes the precipitation. In the second case, the cyclonic circulation around a Low moving along the coast brings in moist air from over the ocean, and the mountains provide additional lift to increase the precipitation. Wet thunderstorms are common, and lightning fires are few.

Heavy snows in the northern coniferous forests persist well into spring. The leafless hardwoods in the areas of lesser snow cover expose the surface litter to drying influences of the sun and strong winds during the spring months. Both the conifers and hardwoods are susceptible to cumulative drying during the fall.

The fire season usually lasts from April through October with peaks in the spring and fall. Drought years are infrequent but may be severe.

The synoptic weather types associated with high fire danger in this region are the Pacific High, Hudson Bay High, Northwest Canadian High, and Bermuda High. All of these types have been described above.

The vegetation in the Great Lakes region consists mainly of aspen, fir, and spruce in the north and some additional hardwoods in the south. There are several upland areas, including the western slopes of the Appalachians, but most of the region has been heavily glaciated. Winter temperatures are quite cold, and summer temperatures are variable. In summer, the region is subjected to cool cP air masses from the north, warm and moist mT air masses from the south, and mild mP air masses from the west.

The annual precipitation in the Great Lakes region is moderate, generally over 30 inches. It is fairly well distributed throughout the year, but most areas have somewhat larger amounts in summer. Winter precipitation is mostly in the form of snow, and the greatest amounts occur with intense cyclones involving mT air masses. Summer precipitation is largely in the form of showers and thunderstorms. Lightning fires are common on both sides of the St. Lawrence and in the northern Great Lakes area.

Strong winds are common with intense storms in fall, winter, and spring, and with squall lines and strong cold fronts in the summer. Humidities are normally moderate to high except during brief periods when cP and mP air masses are warmed by heating and subsidence before much moisture can be added to them.

The Great Lakes are sufficiently large to influence the climate of portions of the region. Near the shores, when the gradient winds are weak, lake breezes can be expected on summer days. The lake breeze is cool and humid and moderates the summer climate along the lake shores.

On a larger scale, the Great Lakes modify air masses that pass over them. Cold air masses passing over the warmer lakes in the fall and winter are warmed and pick up considerable moisture, resulting in heavier precipitation to the lee of the lakes. The amount of moisture picked up depends to a large extent upon the length of the overwater fetch. In spring and summer, warm air masses are cooled as they pass over the cooler waters of the lakes. If the air mass is moist, fog and low clouds form and drift over the leeward shores.

The Great Lakes also affect the synoptic-scale pressure pattern. In spring and early summer when the lakes are relatively cool, they tend to intensify high-pressure areas that pass over them. In fall and winter when the lakes are relatively warm, they tend to deepen Lows that pass over them. On occasion, they will cause a trough of low pressure to hang back as the Low center moves on toward the east. This tends to prolong the cloudiness and precipitation.

The fire season generally lasts from April through October with peaks in the spring and fall. In hardwood areas, the leafless trees in spring expose the surface litter to considerable drying, which increases fire danger. After the lesser vegetation becomes green and hardwoods leaf out, the fire danger decreases. In fall, the lesser vegetation is killed by frost, the hardwoods drop their leaves, and the fire danger again increases.

The synoptic weather patterns producing high fire danger in the Great Lakes region are usually those involving Highs moving into the region from Hudson Bay, Northwest Canada, or the Pacific. Occasionally the region is affected by a Bermuda High type, but this is infrequent and usually occurs during the period when the vegetation is green. The Pacific High type, which was discussed with the Great Plains region, causes more high fire-danger days than any other type.

The Hudson Bay High and Northwest Canadian High types involve cP air masses that move southward or southeastward from their source regions in Canada and on through the Great Lakes region under the influence of a meridional pattern aloft. These air masses are warmed by surface heating and subsidence as they move to lower latitudes. High fire danger is occasionally found in the forward portion of the air mass, if the front preceding it is dry. But the most critical area is usually the western or northwestern portion of the High. By the time this portion of the High reaches a locality, the air mass has been warmed by heating and subsidence, and the humidity becomes low and remains low until either Gulf moisture is brought into the system or the next cold front passes.

The vegetation in the Southern Rocky Mountain region consists of brush and scattered pine at lower elevations, and fir and spruce on higher ridges and plateaus. Many peaks extend above timberline. As in the Northern Rockies, winter temperatures are quite low, and summer temperatures are moderate for the latitude because of the elevation influence.

Precipitation is generally around 10 to 20 inches annually in the valleys and on eastern slopes, and 30 to 40 inches locally at higher elevations on the western slopes. The heavier precipitation at higher elevations is caused by the additional orographic lifting of mP air masses as they are forced across the Rocky Mountains. Most of the precipitation in the winter is in the form of snow. Precipitation is light but not infrequent during the summer, mostly as thunderstorms. These storms cause wildland fires, but ordinarily the burned acreage is small.

There are strong chinook winds with associated warm and dry conditions in the spring and fall on the eastern slopes of the mountains. These winds sometimes bring subsiding air from high levels in the atmosphere down to the surface and produce extremely low humidities.

The fire season normally extends from June or July through September, but earlier or later periods of critical fire weather may be caused by the chinook winds.

The synoptic patterns which produce high fire danger are the ridge aloft and dry cold-front passages. In addition, the pattern producing chinook winds is important on the eastern slopes. In this pattern, the airflow aloft is usually at right angles to the mountain range, while at the surface, a High is located in the Great Basin and a front is found east of the Rockies. In the area between the front and the Rockies the air flows downslope, winds are strong, temperatures are high, and humidities are acutely low.

Heavy pine, fir, and spruce stands dominate the Northern Rocky Mountain region. Many mountain peaks extend above timberline. The portion of this region in Canada includes the Cordilleran Highlands with numerous mountain ranges and dissecting river courses, in addition to the Rocky Mountains. Winter temperatures are quite low, and summer temperatures are moderate.

Annual precipitation ranges from 10 to 20 inches in the valleys to 40 to 60 inches locally in the mountains. Most of the precipitation falls in the winter and spring in the southern portion of this region, while in the northern portion it is fairly well distributed throughout the year, in most years. Winter precipitation is in the form of snow. In the southern portion, there often is widespread rainfall until June, followed by generally light precipitation during the summer.

There is a gradual drying out of forest fuels during July and August with increasing fire danger. Frequent thunderstorms may occur then but little or no precipitation reaches the surface, so that frequent and severe lightning fires occur in both the Canadian and United States portions of the region. Also, extremely low humidities can result from large-scale subsidence of air from very high levels in the atmosphere. Occasional chinook winds on the east slope of the Rockies produce moderate temperatures and are effective in bringing subsiding air to the surface. The fire season usually extends from June or July through September.

The synoptic weather types producing high fire danger are similar to those described for the Great Basin region. Particularly important are the ridge-aloft pattern which produces warm, dry weather and the patterns producing high-level thunderstorms.

The vegetation in the Southwest (including Sonora, Mexico) is mostly grass, sage, chaparral, and ponderosa pine. The region in which wildfire is a problem is essentially a plateau, but it also includes the southern portion of the Cordilleran Highlands. The low-elevation areas of the Southwest have a large annual range and a large diurnal range of temperatures, the latter being larger in the summer than in the winter. The higher elevations have both lower mean and lower maximum temperatures. Spring and early-summer temperatures are very high during the daytime because of clear skies and low humidities. The extreme southwest low-elevation portions have extremely hot and dry summers, while the higher elevations of the rest of the region have more moderate temperatures, and frequent summer thunderstorms during July, August, and September.

The Southwest is quite dry, with annual precipitation in some areas as little as 5 to 10 inches. This occurs as winter rain or snow, and as rains accompanying the frequent summer thunderstorms. In the first scattered storms in the late spring and early summer, little precipitation reaches the ground. Later in the summer, thunderstorms are usually wet.

The Southwest is characterized by an annual minimum of fire danger in the winter months and a secondary minimum in August. The most dangerous fire season is generally May and June

Since the Southwest has a generally high level of fire danger in spring and again in fall, the important synoptic patterns are those which cause peaks in fire danger or those which cause dry thunderstorms. The most critical fire weather occurs with a broadscale pattern aloft showing a ridge to the east and a trough to the west of the region, and southwesterly flow over the region. The fire danger peaks as short-wave troughs move through this pattern and cause a temporary increase in wind speed. The pattern favorable for thunderstorms has the subtropical High aloft to the north of the region, and southeasterly flow bringing moist air from over the Gulf of Mexico to the Southwest region. When this pattern becomes established, the first moisture brought in is usually in a shallow layer aloft. The resulting thunderstorms tend to be of the dry, fire-starting type and appear when the fuels are dry and the fire potential is high. Then, as the pattern persists, moisture is brought from the Gulf of Mexico in a deep layer, and the thunderstorms produce rain which reaches the ground and reduces the fire danger.

Fire behavior describes how a fire burns- where it burns, how fast it travels, how much heat it releases, and how much fuel it consumes. It is important to understand what controls fire behavior and how to predict it. This knowledge will help predict fire effects, conduct prescribed burns, predict wildfire risk, and control wildfires. The following sections provide a basic understanding of what controls fire behavior, how to describe fire behavior, and how to predict fire behavior.

• Combustion and Heat Transfer explains how fire produces heat through the physical-chemical process of combustion and how this heat is transferred through pathways such as convection, radiation, and conduction.
• Fire Spread defines the four primary modes of fire spread: ground fires, surface fires, crown fires, and spotting.
• Fire Growth explains how a fires shape and rate of spread changes in relation to fuel, weather, and topography.
• Extreme Fire Behavior explains causes and characteristics of extreme fire behavior and describes fire whirls, plume-dominated fires, and vortex phenomena.
• Flame Characteristics defines and explains how to measure flame height, length, depth, angle, and char height.
• Fire Behavior Prediction Systems briefly describes various computer models available for predicting fire behavior based on in fuels, weather, and topography. This section also presents fuel models.
• Fire Danger Rating Systems provides background of the various qualitative and numeric indices of fire potential used to estimate present and future fire danger for a given area, such as the National Fire Danger Rating System.

Fire releases heat through combustion. Combustion is a physical and chemical process that unleashes the solar energy stored in chemical form in fuels as heat. Oxygen, heat, and fuel often called the fire triangle must be present in the proper combination for a fire to ignite and initiate combustion. Sufficient oxygen is found in our atmosphere, fuel accumulates in forested ecosystems so when it is dry enough all that is required is the heat or ignition source, which can be either from a natural or anthropogenic source.

A common sequence of physical processes occurs in all fuels before the energy contained in them can be released and transferred during combustion. The changes a fire goes through are traditionally organized into five phases of combustion: pre-ignition, flaming combustion, smoldering combustion, glowing combustion, and extinction. Although flames are the most recognizable and significant phase of the combustion process, these other phases have important impacts on ecosystem resources as well.

For a fire to grow and spread, heat must be transferred to surrounding fuel. Heat produced during combustion is transferred to other fuels, the soil, and the atmosphere primarily by convection, radiation, and conduction. As wildland fires spread, radiation and convection are the primary methods of heat transfer from one fuel to another, while conduction transfers heat within large pieces of fuel. Mass transport by spotting or fuels rolling downslope can also be an important means of heat transfer during intense fires. When water is present, heat transfer can also occur by vaporization. Fire intensity is an important measure of the rate of heat released by a fire.

Also see Heat Transfer.

Fire Triangle

The fire triangle illustrates the relationship between the three basic elements needed to establish fire: oxygen (O₂), combustible fuel (a form of (C₆H₁₀O₅)_n), and heat to initiate and sustain combustion. These three elements combined together in the proper ratio lead to fuel ignition and combustion. The progression of combustion in wildland fires is generally divided into four phases: pre-ignition, flaming, smoldering, glowing, and extinction. Here, the chemical reaction of combustion is explained.

In the combustion process, oxygen “oxidizes” (combines with) carbon and hydrogen molecules as heat breaks chemical bonds in carbon chains during combustion. When vegetation is burned, heat is released in the form of thermal, radiant, and kinetic energy.

Heat is supplied initially from an external source (lightning, matches, etc.) that begins breaking chemical bonds in the organic compounds. Ignition occurs when the external heat source has broken enough carbon bonds and released enough heat from these bonds to sustain combustion without the external heat source. Thereafter, heat generated by the breaking of bonds maintains combustion as a “feedback loop” until the combustible fuel is exhausted or removed, or the heat produced is inadequate to further break additional bonds, or oxygen supply becomes limiting.

The general reaction formula for combustion is the reverse of photosynthesis, and is represented as:

(C₆H₁₀O₅)_n = a plant-synthesized organic compound

O₂ = gaseous oxygen

CO₂ = carbon dioxide

H₂O = water vapor

PICCs = products of incomplete combustion (multi-carbon molecules pyrolyzed from longer compounds)

Combustion can be high-speed (as in rocket motors, gas turbines or internal combustion engines) or low-speed (forest fires and candles). In low-speed combustion, a feedback loop exists between gases above the fuel and the fuel itself. In the gaseous phase, oxygen in the air drives combustion of the gaseous fuels and heat is released exothermically. Some of the heat sustains combustion of the gases and some is transferred back to the condensed phase (vegetative fuels), where it causes endothermic (heat/energy absorbing) gasification of the vegetative fuel. The feedback of heat from gas phase flames to condensed phase fuel is a crucial mechanism for sustaining the combustion process.

With the heat losses, combustion of plant material in open air is seldom complete; if it were, all carbon molecules would be oxidized to CO₂. Products of incomplete combustion include carbon monoxide (CO), nitrous oxides (NOx), sulphur oxides (SOx), dioxin, a variety of multi-carbon molecules and soot (Ward 2001).

Wildland fuels vary in their combustibility depending on type, chemical composition, size, arrangement, moisture content, and quantity. Vegetative fuels are generally composed of cellulose, hemicellulose, other sugars, lignin, resins, and other organic compounds (see Chemical Fuel Properties). Dead plant material generally contains less moisture (due to a cessation of water uptake), a different arrangement of moisture (less intra-cellular water), and less long-chain organic compounds (due to decomposition), all of which influence combustibility (or fuel availability).

Wildland fires can be caused by both natural and anthropogenic mechanisms. While types of ignition are relatively standard across the southeastern U.S., the rates at which certain sources occur are not. For example, in Florida the most common sources of ignition from 1998 to 2002 were lightning (25%), incendiaries (20%), debris burning (18%), and equipment (Florida Division of Forestry 2003; see also Florida Forest Protection Bureau). In North Carolina the three most common sources of ignition from 1997 to 2001 were debris burning (41%), incendiaries (20%), and children (9%) (North Carolina Division of Forest Resources 2001).

Lightning ignition

“Spontaneous” ignition

Wildfires ignited by humans can be either accidental or deliberate. Nationally, a major source of accidental human-caused wildland fire is burning yard waste or trash (debris burning) that becomes uncontrollable. Other major sources of accidental ignition include escaped campfires, discarded cigarettes, fireworks, sparks from railroad equipment, and highway/off-road vehicles. Railroad equipment may cause fires in a variety of ways: improperly maintained tracks may cause sparks during transit (especially on steep grades); maintenance equipment (such as cutting and grinding tools) may cause sparks; brake pads may shed embers and sparks in periods of high stress. Vehicles, either highway or off-road, cause fires through hot exhaust systems, either through contact with dry vegetation (when parked) or through the production of sparks during normal operation (some vehicles have been outfitted with spark-arrestors to reduce this risk) (Baxter 2002). Most accidental fires caused by discarded cigarettes, railroad equipment, and vehicles occur in extremely dry, low humidity situations–a common condition in the Southeastern U.S. during dry spring and fall periods.

Deliberate anthropogenic fires fall into two major categories: arson and prescribed fires.

• Arson or incendiary fires are intentionally ignited utilizing a variety of methods: dropping incendiary devices (matches, flares, etc) or arranging simple solar heating systems. Approximately 25% of wildfires are intentionally set in the United States.

• Prescribed fires are used in the Southeast to meet specific land management objectives and to maintain natural ecosystems. Qualified land managers set these fires under specific conditions of humidity, wind, and fuel moisture. Prescribed fire is applied to more acreage of land in the Southeast than in any other region of the United States.

• Pre-ignition is the heat absorbing phase of combustion where heat is applied to fuel resulting in vaporization of water and other substances providing the gases that sustain flames in the next phase of combustion.

• Flaming combustion is a phase of combustion where heat is released. This phase is the most efficient phase of combustion producing the least amount of smoke per unit of fuel consumed. The products of flaming combustion are primarily carbon dioxide and water vapor.

• Smoldering combustion is the least efficient phase of combustion and produces the most smoke. This phase lacks flame, and is associated with conditions where oxygen is limited – either by char of fuels (particularly those with large surface to volume ratios) or by tightly packed fuels like duff and organic soils or in wet fuels.

• Glowing combustion is the phase of combustion when only embers are visible. During this phase, there is no longer enough energy to create visible smoke.

• Extinction either occurs when fire goes out after the available fuel is consumed or there is insufficient heat produced by oxidation in previous stages to vaporize moisture from unburned fuel. The fuel moisture, surrounding air or inorganic materials absorb and reduce the heat from combustion and this results in the fire extinction process (DeBano et al. 1998).

Pre-ignition is the heat absorbing (endothermic) phase of combustion where heat is applied to fuel resulting in vaporization of water and other substances providing the gases that sustain flames in the next phase of combustion, flaming combustion. During pre-ignition, the temperature of the fuel is raised to at least 100° C by radiation, convection and/or conduction. Heat of pre-ignition is the amount of heat required to raise fuel to ignition temperature. The first major product of pre-ignition fuel heating is water vapor converted from liquid water contained within cells and loosely tied to cell wall structure. Fuel temperatures will not rise above 100° C until this water is volatilized. During this phase, low temperature waxes, oils and other extractives are also degraded, yielding volatile gases such as lipids and other highly combustible hydrocarbons. This process is called pyrolysis.

Pyrolysis is the thermal degradation of chemicals in which the bonds that hold complex molecules together are broken, reducing molecule size and releasing smaller molecules from the material as gases (Pyne et al. 1996). It generally occurs in an oxygen-deficient environment. Pyrolysis occurs at different rates for different compounds. Pyrolytic degradation of polymers such as cellulose is exothermic in the presence of oxygen and endothermic in the absence of oxygen. Both processes produce char that can react with oxygen in an exothermic reaction. Char is the “black solid residue of a variable aromatic nature remaining after the end of the initial rapid weight loss upon heating of a polymer” (Ohlemiller 1985). Volatilization involves the pyrolysis of large compounds in the fuels to a gaseous state. Some resins, waxes and oils of woody materials volatilize to a gaseous state at relatively low temperatures. These volatiles are often released in the highest concentrations during the pre-ignition phase.

Pyrolysis temperatures vary for different organic compounds; hemicellulose begins degradation at 200 to 260^oC, cellulose at 240 to 350^oC, and lignin at 280 to 500^oC (Roberts 1970). Pyrolytic decomposition of cellulose also varies in different temperature conditions. At low temperatures pyrolysis of cellulose produces a variety of products including char, carbon dioxide (CO₂), carbon monoxide (CO), methane (CH₄), hydrogen (H₂), ethane C₂H₄, water, and low-molecular weight acids (Shafizadeh and DeGroot 1976); higher temperatures (280^oC and above) produce levoglucosan, a volatile fuel that supports a gas-phase flame.

Flaming combustion is the most efficient phase of combustion producing the least amount of smoke per unit of fuel consumed. The transition from pre-ignition to flaming combustion is the transition of the overall combustion reaction from a heat absorbing (endothermic) to a heat releasing (exothermic) reaction and occurs at fuel temperatures just above 300^oC. Flaming combustion is initiated by the actual ignition of the volatile gases produced during the pre-ignition phase and located above the heated fuel. A principle characteristic of the flame is that the fuel (gases) and oxidizer (air) are initially separate and combustion occurs in the zone where the gases mix. Flames are not attached directly to the surface but are separated by a thin layer of vapor or gas. In the flame, thermal degradation of molecules occurs more completely and rapidly than in pre-ignition, with large releases of heat energy. This heat induces further volatilization of organic compounds in the solid fuel, which sustain the flame. Ignition tends to occur as a chain reaction, with the ignition of fuels at a particular location initiating the ignition of adjacent fuels. A line of fire can be considered a chain of ignition events.

The energy that maintains the flame and chain reaction of ignitions is known as the heat of combustion, which is the amount of heat released when a unit of fuel is oxidized completely. The heat of combustion for most wildland fuels is approximately 18,620 KJ/kg (8000 Btu/lb). Temperatures reached during flaming combustion (after ignition) can vary dramatically depending on the amount of fuels and weather conditions. Very intense fires have been recorded as reaching temperatures of 1650^oC, but average temperatures tend to fall within the range of 700-980^oC (Pyne et al. 1996). In the flaming phase, combustion efficiency is relatively high and emits the least quantity of pollutants compared with the mass of fuel consumed. Combustion efficiency is the fraction of fuel carbon converted to carbon as CO₂ (USDA 2002). For more information see Flame chemistry.

Products of flaming combustion

The products of flaming combustion are predominately carbon dioxide (CO₂) and water vapor. This water vapor is not the result of the heating of the fuels as in the pre-ignition phase but rather a result of the oxidation of organic molecules. Other combustion products contained in the smoke above a flame include small solid particles and multi-carbon molecules of various lengths and structures that are not fully broken down into basic carbon molecules (i.e., CO and CO₂) while passing through the flame, but are light enough to be lifted by convection above the flames. Some organic gases cool and condense without passing through the flame zone. Others pass through the flames and are only partially oxidized producing a great variety of emissions. Some compounds with high molecular weights cool and condense into tar droplets and soot (solid particulates). This particulate matter makes up part of the visible smoke, and affects ambient air quality and human health. Particles can be separated into a fine-particle mode with a mean diameter of 0.3 micrometers and a coarse particle mode with mean diameter of 10 or more micrometers. The percentage of fine particles produced ranges from 80 to 95% depending on the turbulence in the combustion zone. Fires of low intensity, where fire is barely sustained and inefficient in combustion, increase particulate matter emissions (Ward 1998).

Flaming combustion tends to follow a cycle beginning as an intense fire that causes a high rate of volatilization and ignition of surrounding materials. As the fire progresses, combustion levels off, and as the amount of fuel available for combustion decreases and the amount of carbon and ash on the surfaces of the fuels increases, combustion decreases. When the amount of volatiles released by the fuels is so reduced that it can no longer maintain a flame above the fuel surface, combustion enters the smoldering phase. Consumption of plant material usually ranges from 50-95%, depending on fuel characteristics. Dry fine fuels, such as cured grasses, tend to be more completely consumed during combustion than moist heavy fuels (see Fuel Availability and Consumption).

Extractives and other non-cellulosic compounds have different requirements for combustion, and can emit products such as hydrocarbons, alcohols, aldehydes, methane (CH₄), nitrous oxides, and sulfur oxides. Sulfur emissions such as carbonyl sulfide are essential in the synthesis of plant amino acids. Methyl chloride is produced in greater quantities during the smoldering phase than the flaming phase (inversely proportionate to the rate of heat release). Methane and other hydrocarbons are also produced in greater quantities during the smoldering phase. Products of incomplete combustion contribute to air pollution and are a concern for land resource managers.

Flames are a phenomenon associated with the chemical reaction of oxygen with flammable gases that have been heated above their ignition temperature. The shape of the flame is commonly that of a hollow cone, with two distinct zones: an inner, cone-shaped area consisting of unburned gas and little oxygen, and an outer visible area in which chemical reaction associated with flaming combustion occur (pyrolysis with the formation of CO₂ and H₂O).

Typically flames are described as either premixed or diffusion flames. Premixed flames are those in which the fuel and oxidizer are mixed before burning, as in a Bunsen burner, and the fuel-oxygen ratio is constant everywhere in the flame. In diffusion flames (candles or forest fires), the fuel and oxygen are separated before burning, and the fuel-oxygen ratio varies throughout the flame. Diffusion and premixed flames are also classified as laminar, turbulent, or transitional. Laminar flames have smooth, steady flow characteristics in and around the flame. Turbulent flames have irregular, disorganized flow characteristics in and around the flame caused by variations in fuel components and mixtures and the way in which they burn. Some flames may be a mixture of laminar and turbulent.

Combustion of vegetative fuels is a diffusion flame process. Initial heating of unburned fuel releases volatile molecules and begins to fragment long-chain cellulose and lignin polymers into a variety of organic compounds, with many released in a gaseous state. As the gaseous fuel mixture migrates from the solid fuel surface through the oxygen-deficient flame interior, and toward the flame front, it is heated by radiation from the flame zone upstream. At the same time, oxygen from the atmosphere diffuses inward. Once enough oxygen is added to the mixture and temperatures are hot enough to sustain combustion, the mixture becomes incorporated into the flaming reaction zone. It is difficult to draw a distinct line between preheat and reaction zones but it can be thought of as the point at which exothermic reactions become significant. In the reaction zone around the flame interior, oxygen and the gaseous fuels mix in proportions that support the rapid combustion/oxidation chemical reactions that result in the visible emission of light (flame). The heat from this flame structure and further diffusion and turbulent mixing of oxygen at the surface of the solid fuel help sustain the pyrolytic decomposition of cellulose, lignin, and other solid fuel hydrocarbons. Gases other than products of complete combustion that emerge from the flame enter the burned gas zone where products of incomplete combustion may reaggregate or solidify as they are carried away from the flame by convection. A diffusion flame zone is substantially thicker than premixed flames because of the process of mass diffusion of fresh gaseous fuel mixture into the flame zone. Once carbon and ash begin to build up on the solid fuel surface, the pyrolytic reactions no longer produce sufficient fuel gases to maintain the flame envelope.

The light emitted by flames is characteristic of the specific chemical reactions that are taking place. The luminosity of the flame is usually caused by solid particles present in the burning gas and heated to incandescence. The distinctive yellow color results from the formation of soot (solid carbon particulates). Combusting fuel gases burn with a blue color. The orange color of flames is due to the radiation from an abundance of small solid particles. The emission spectrum is continuous, and the color of the flame depends on the concentration of these particles.

Smoldering combustion is the least efficient phase of combustion and produces the most smoke per unit of fuel consumed. This phase lacks flame, and is associated with conditions where oxygen is limited – either by char of fuels (particularly those with large surface to volume ratios) or by tightly packed fuels like duff and organic soils or in wet fuels. As such, smoldering combustion dominates during ground fires. The smoldering phase can be distinguished from glowing combustion (the final phase of combustion) by the presence of smoke.

Smoldering combustion is “a self-sustaining, low temperature combustion process involving pyrolysis of the substrate ahead of a solid-phase combustion front” (Shafizadeh et al. 1982). The main features include thermal degradation and charring of the fuel with evolution of smoke (Moussa et al. 1976). A build-up of char on the surface of fuels is usually the reason for the transition from flaming combustion to the smoldering phase. Char is residual carbon that does not fully break down from pyrolysis, and therefore builds up on the surface of fuels. If the levels of char that inhibit volatile release are reduced in particular regions, fuels in the smoldering phase can re-ignite.

Smoldering fire can burn in duff and other ground fuels for long periods of time after flames have ended, and can be responsible for high levels of fuel degradation. Smoldering also tends to occur in tightly packed materials that lack the high levels of oxygen needed to maintain flaming conditions, or in fuels that have high moisture content and therefore do not readily ignite. Surface temperatures of fuels in the smoldering phase can exceed 500^oC due to the energy release from pyrolysis in the fuel (Pyne et al. 1996). The smoldering process is important in ecosystems that require removal of duff for plant regeneration.

The main factors that influence the smoldering rate of spread are time of day, wood chemical and physical composition, and log or duff location on the ground (Rabelo et al. 2004). Moisture content seems to play a less significant role in the speed of propagation than other combined parameters. Fire burning in high drought conditions can be self-sustaining and extend deep into the duff. Smoldering combustion varies with fuel type; it is more prevalent in duff, organic soil, and rotten logs, and less in fuel with high surface to volume ratios, like grasses, shrubs, and small diameter woody fuel (Sandberg and Dost 1990). See also: Ground Fires.

Products of smoldering combustion

Glowing combustion is the phase of combustion when only embers are visible. Glowing combustion refers to the process of surface oxidation of solids occurring in the final stage of the smoldering process (Simmons 1995). The glowing phase of combustion occurs when there is no longer enough energy to create the aerosol smoke characteristic of the smoldering phase, and thus there is no longer a release of tars, volatiles, or liquids from the fuels. Oxygen in the air can now reach the fuel surface and the charcoal begins to burn with a characteristic yellow glow with no visible smoke.

The main products released by fuels in the glowing phase are the invisible gases carbon monoxide (CO) and carbon dioxide (CO₂). Although the temperatures associated with the glowing phase are usually around 600ºC, much lower than the temperatures associated with flaming combustion, the heat remains in a particular location for longer duration thus large amounts of organic decomposition can occur during this phase (Pyne et al. 1996). The glowing phase eventually ends when the temperature of the surrounding char drops to a point where pyrolysis can no longer occur.

Heat produced during combustion is transferred to other fuels, the soil, and the atmosphere primarily by convection, radiation, and conduction. As wildland fires spread, radiation and convection are the primary methods of heat transfer from one fuel to another, while conduction transfers heat within large pieces of fuel. Mass transport by spotting or fuels rolling downslope can also be an important means of heat transfer during intense fires. When water is present, heat transfer can also occur by vaporization.

Conduction is the transfer of heat energy by the motion of adjacent molecules, from a region of high temperature to a region of lower temperature. It occurs in all solid materials, with greatest efficiency in dense solids such as metal. Dense materials conduct heat better because the molecules are much closer together. Wood, soil, and water are considered poor conductors, while material such as steel or bronze readily conduct heat. A dense material conducts heat better into the interior of the substance so the surface heats up more slowly. This explains why solid, dense wood requires a lot of heat to ignite, but rotten wood can be ignited with a spark. Since the wildland fire environment is generally devoid of dense materials (except logs and tree stems) conduction is usually not a primary method of heat transmission except during the combustion of logs, duff, soil organic matter, and roots (Countryman 1976).

Conduction is expressed by the equation:

q”_x = -k (DT/Dx)

where q”_x = heat transferred per unit time, k is a constant that represents thermal conductivity of a particular material, DT is temperature change over time, and Dx is change in distance over time.

Convection is the transfer of heat from one point to another by movement of the surrounding fluid or gas. In the case of wildland fire, convection is exclusively the movement of heated air. “Free convection” occurs when the fluid motion of the liquids or gases creates convection currents because of differences in densities caused by temperature differences of two similar substances.

Wind acts as a force that drives convection, and the stronger the winds are the further away fuels are pre-heated by convection. While convection currents always rise (as does all heated air), they may be directed laterally by wind, and may be trapped (slowed in their dissipation) by thermoclines (air stratification along thermal levels). See also: Convective Winds.

Convection is the primary mode of heat transfer to shrubs and other tall vegetation in wildland fires. Rising convective currents influence the spread of fires from the understory and midstory to tree canopies. In high-wind situations and on slopes, convection may carry heat into adjacent fuels, greatly increasing the rate of fire spread. Convection is also critical in its ability to carry firebrands (sparks or embers) up and out from a fire front and depositing them outside the fire’s perimeter, causing spot fires. In general, convection currents can be tracked by the flow of particulates from a flame front (smoke).

Convection is described by the equation:

q”_x = h DT

where, h is a non-constant coefficient that varies under different conditions and DT is temperature change over time.

Radiation is the transfer of heat by electromagnetic waves. These waves travel at the speed of light away from the heat source and transfer their energy to matter they impact (the denser and more opaque the matter, the greater the energy transfer rate). Unlike conduction, the heat of one body does not have to be in contact with another body to transfer heat by radiation.

Radiation plays a key role in the preheating of fuels in forest fires, especially in the absence of wind, where heat transfer by convection would be reduced. Radiation travels in straight lines so intensity depends on the angle of the objects absorbing the electromagnetic waves. A perpendicular angle will yield the most intense heat transfer. This is also important for fuel arrangements and the effect of slopes on wildland fire.

The intensity of radiant energy can be summed up in the equation:

q”_x = tsT⁴

where t is the emissivity, a measure of how well the surface performs as a radiator, s is the Stefan-Boltzmann constant (5.67 x 10^-8 W/m² K^-4), and T equals temperature in degrees Kelvin. For a point source of radiation, the radiation intensity decreases inversely with the square of the distance from a specific point. This means that the radiation intensity 10m from the source is only ¼ that at 5m. As the distance from the source increases, the same total amount of radiation is spread over a greater area and so the amount of heat received per unit area decreases.

Fire intensity is a measure of the rate of heat released by a fire. It includes both radiant and convectional heat. There are several definitions and ways to measure fire intensity. The most common of these is fireline intensity. Other measures of fire intensity include reaction intensity, radiant intensity, convection intensity, total fire intensity (defined below).

Fire intensity is important for planning prescribed burns because it directly influences the smoke column height and flame heights during a burn, as well as the responses of an ecosystem to a particular burn (e.g., fire effects). Fire intensity is also useful in assessing the difficulty of wildfire containment (see Fireline Intensity).

Fire intensity is directly proportional to a fuels heat of combustion, the amount of fuel consumed, and a fires rate of spread. As such, fuels, weather, and topography are very important in determining the rate of heat released by a fire. Fuel properties that directly or indirectly affect fire intensity include fuel loading, moisture content, arrangement, chemical composition, and size. Wind speed and other weather conditions that influence fuel moisture also influence fire intensity. Exactly how these fuel, weather, and topographical conditions affect fire intensity and heat release is explained in the following pages:

• Effects of weather on fire intensity and rate of spread
• Effects of fuel on fire intensity and rate of spread
• Effects of topography on fire intensity and rate of spread

Measures of fire intensity

• Fireline intensity (also known as Byrams fireline intensity or frontal fire intensity) is the rate of heat energy released per unit time per unit length of fire front, regardless of the depth of the flame zone (Byram 1959).
• Reaction intensity (also called combustion rate or area-fire intensity) is the rate of heat release per unit area per unit time in the flame zone. Reaction intensity can be derived by dividing fireline intensity (kW/m) by flame depth (m). Under laboratory conditions, reaction intensity can be measured by recording weight loss and determining heat release through calorimetry. In the field, reaction intensity is calculated by estimating the amount of of fuel burned per second and assuming heat yields for the fuel (Chandler 1983, DeBano et al. 1998). It is usually expressed kW/m². Reaction intensity is calculated in Rothermels (1972) fire spread model which, in turn, is used for fire behavior calculations in the National Fire Danger Rating system (NFDRS) and BEHAVE (Wade 1986).
• Radiant intensity is the rate of thermal radiation emission either across the entire radiomagnetic spectrum or within specified wavelengths (spectral intensity). Radiant intensity is intercepted at or near the ground surface, or at some specified distance ahead of the flame front (Chandler et al. 1983, DeBano et al. 1998). The units are usually expressed in cal/cm²sec.
• Convection intensity is the heat from a fire that lifts gases and entrains air above the flame zone. Convective intensity is not determined directly, but is the difference between the total energy release rate and the sum of conductive and radiant heat losses (Chandler et al. 1983, DeBano et al. 1998). Convective intensity calculations are used to predict convection column phenomena. The units are usually expressed as kcal/m²min or as kW/m².
• Total fire intensity, the rate of heat output of the fire as a whole, is a function of the rate of area burned, fuel loading, and estimated heat yield. Its not useful in describing site-specific effects (Chandler et al. 1983, DeBano et al. 1998).

Fireline intensity, also known as Byrams fireline intensity or frontal fire intensity, is the rate of heat energy released per unit time per unit length of fire front, regardless of the depth of the flame zone (Byram 1959). It is calculated as the product of available fuel energy and the fires rate of advance:

I = Hwr

where:

• I = Byrams fireline intensity (kW/m or Btu/s/ft)
• H = fuel low heat of combustion (kJ/kg), subject to reductions
• w = weight of fuel consumed per unit area in the active flaming zone (kg/m2)
• r = rate of spread (m/s)

Fireline intensity may vary by more than 1000-fold, or from 15 to at least 100,000 kW/m (Alexander 1982). However, most fire intensities seldom exceed 50,000 kW/m and most crown fires fall within the range of 10,000-30,000 kW/m. Low intensity surface fires are generally less than 550 kW/m. Fireline intensities above 4000 kW/m are generally characterized as "high intensity" (Alexander 1982). Because considerable burning can take place after passage of the flaming front, these figures do not describe the total energy or heat released (Alexander 1982).

The wide variation in fireline intensity is largely because of the potential variation in rate of spread. Rates of spread can vary 100-fold, while fuel consumption varies only 10-fold, and low heat of combustion varies so little from fuel to fuel (roughly 10%) that it can be thought of as a constant (Alexander 1982). Because of the importance of rate of spread, fireline intensity is greatly influenced by weather and topography (Brown and Davis 1973, DeBano et al. 1998).

While fireline intensity may be one of the best single descriptors of fire behavior (Alexander 1982), it is difficult to accurately measure the required variables in the field. In particular, it is difficult to distinguish fuel consumed in the active flaming zone (w) from fuel consumed after the flame front passes. As a consequece, w is often overestimated (Wade 1986). An easier way to estimate fireline intensity in the field is to determine fireline intensity using observed flame length. Although not entirely accurate, estimating fireline intensity indirectly from flame length will help avoid errors due to overestimating w when using Byrams equation. A more complete description of fire behavior would quote the rate of spread in addition to fireline intensity because identical fireline intensities can be arrived at by fires with different spread rates and fuel consumption (Van Wagner 1965, Alexander 1982).

Fireline intensity can be useful for comparing fires (Wade 1986). It can also be used to assess the effects of prescribed burns. For example, fireline intensity correlates well with crown damage and can actually be used to calculate lethal scorch height as well as expected temperatures at certain heights above surface fires. Other fire parameters are better for assessing damage to plants within and below the flame zone, however. For example, residence time or reaction intensity are more strongly correlated with basal stem damage than fireline intensity. Depth of burn or heat release per unit area are good indicators of damage to roots and stems belowground (Wade 1986).

Fireline intensity is also used to assess the difficulty of wildfire containment. For example, direct attack with hand tools and assured control of prescribed fires is possible when fireline intensity is less than 400-425 kW/m (Hodgson 1968, Chandler et al. 1983). Heavy mechanical equipment can usually control a fire if fireline intensity is below 1700 to 1750 kW/m (Forest Service 1978, Chandler et al. 1983). Spot fires can become serious at 2000 to 2100 kW/m (Hodgson 1968, Chandler et al. 1983) and fires are completely uncontrollable with fireline intensities above 3500-3700 kW/m (Forest Service 1978, Chandler et al. 1983).

Estimating flame length from fireline intensity

The following equations can be used to estimate fireline intensity from flame length:

where:

• L = flame length (m)
• I = fireline intensity (kW/m)

The last equation is often simplified for field use as:

The results of this simplified equation are within 20% accuracy of actual fireline intensity (Chandler et al. 1983), which is generally adequate considering it is difficult to estimate flame length to better than 20% accuracy (DeBano and others 1998). Numerous observations of flame length (minimum, maximum, and mode- or most frequently occurring flame length) should be considered with these equations (Wade 1986).

Estimating flame length from fireline intensity

Relationship between fireline intensity, ambient temperature, and scorch height

The following equations can be used to estimate flame length from fireline intensity:

where:

• L = flame length (m)
• I = fireline intensity (kW/m)

For crown fires, half of the mean canopy height should be added to L.

Relationship between fireline intensity,
ambient temperature, and scorch height

Van Wagner (1973) developed the following graph depicting the relationship between fireline intensity, ambient temperature, and crown scorch height base on the below equation using Canadian conifer species. This graph is commonly used prior to ignition to calculate the scorch line from the expected weather conditions and fire behavior. It likely overpredicts scorch height in southern pines (Wade 1986), so it will err on the safe side if used to solve for scorch height.

Comparing this predicted height with the height of the stand to be burned can help project expected crown scorch. Although solving the equation below for I provides a way to reconstruct fireline intensity after burns using crown scorch, Cain (1984) reported that this technique seriously underestimates fireline intensity in southern pines.

Equation for estimating crown scorch height from fireline intensity

S = (35/ (60 - T)) * ( I^7/6/(0.79(I +47W³)^1/2))

where:

• S = scorch height (m)
• T = air temperature (C)
• I = fireline intensity (kW/m)
• W = windspeed (m/sec) (Van Wagner 1973)

Expected temperatures at any height above a surface fire can be estimated or calculated from ambient temperature and fireline intensity (Van Wagner 1973,1975):

where:

• T = temperature rise above ambient (C)
• I = fireline intensity (kW/m)
• h = height above ground (m)

Fuel Type and Wind Effects on Rate of Spread

Wind speed has one of the greatest effects on fire intensity and rates of spread. As wind blows across a fire, it pushes the flame forward and closer to the unburned fuel in front of the fire (see Flame Tilt). This increases convection and radiation, which dry the fuel and increase its combustibility. In general, the higher the wind speed, the further the flame leans and the faster it dries the fuels, increasing both fire intensity and rate of spread. Wind also adds oxygen to the existing fire, further increasing combustion rates in the flaming zone (Whelan 1995).

The following equation illustrates the direct relationship between fireline intensity (I, kw/m) and wind speed (U, m/s): I = 17.5U³ (Roberts 1976). This relationship has also been quantified in fire behavior prediction models such as BEHAVE, and therefore can also be seen by running fire behavior simulations under different wind conditions (see: Heat release rates under different fuel and weather conditions). The graph below similarly illustrates the effect of windspeed on rates of spread.

Wind also influences the direction of spread and can carry sparks and firebrands downwind of fires, greatly increasing spread rates (see Spotting). A shift in wind direction could rapidly turn a slower moving flanking fire to a head fire, increasing its rate of spread.

Weather conditions that control fuel moisture contents, such as temperature, relative humidity, and precipitation, are also important since fuel moisture determines fuel availability. Weather variations over a given time period generally have a greater effect on fire intensity than fuel variations in the stand. Weather changes can occur in a matter of minutes, whereas fuel variations take place over longer periods of time (Bessie and Johnson 1995).

The following table depicts how different fuel and weather conditions influence heat release and fireline intensity. Table: Fuel type, wind speed, and fuel moisture effects on fireline intensity compares total heat release and fireline intensity for three fuel classes under various wind speed and fuel moisture conditions. The results were generated by the Behave Plus 2.0.2 fire behavior simulator (Andrews and Bevins 2003). The fuel models for tall grass (FM 3), southern rough (FM 7), and light logging slash (FM 11) were used to represent typical fine, medium, and large fuels found in the southeast, respectively (Anderson 1982). These fuel models are summarized in (Table: Fuel Model Parameters).

For all fuels, the highest intensity fire (highest rate of heat release) occurs under conditions of high wind speed and low moisture content. Tall grass burns at a much higher intensity than the other fuels under the same weather conditions, due to the high surface area-to-volume ratio of grass, which allows it to lose moisture more rapidly than larger fuels, the higher loading of 1-hour (<.25 in. diameter) fuels, and the faster rate of spread in the grass fuels. On the other extreme, large fuels will not burn at 18% moisture content because they are unable to efficiently lose water. Total heat release per unit area of fuel does not change with wind speed but decreases with increased moisture content, because there is less available fuel per unit area as fuel moisture content increases. Another important consideration in interpreting the table is that fine fuel consumption is by flaming combustion whereas large fuels are consumed by glowing and smoldering combustion.

For more information on the effects of fuel and weather on heat release and fire intensity, detailed physical and thermochemical process equations can be found in Johnson and Miyanishi (2001) and Drysdale (1999).

Slope steepness, elevation, aspect, and configuration of land are all variations in topography that influence fire intensity and rate of spread.

Slope steepness

Slope steepness affects fire behavior in a similar way as wind by affecting the flame angle. As slopes increase, flames are closer to fuels upslope from the flaming front. This preheats fuel, increasing its combustibility and rate of spread (Rothermel 1985). The effects of slope on rate of spread and flame length for a typical open pine stand with a grass understory are illustrated in the following figures. Both parameters, as modeled by Behave Plus 2.0, increase at a faster rate as slopes increase from zero to 50%.

Slope effects on rate of spread

Slope effects on flame length

Elevation

Elevation influences temperature, precipitation, and relative humidity, all of which affect fuel moisture and in turn, fuel availability. As elevations increase, temperatures tend to decrease while precipitation and relative humidity increase; all these changes reduce rate of spread. One exception to this general pattern is the occurrence of thermal belts at midelevations where warm air at night is trapped below cooler air at both higher and lower elevations.

Aspect

Aspect is the direction a slope is facing, which affects fire behavior due to differences in wind and solar radiation. Solar radiation, in turn, affects relative humidity and fuel types. Ignition and spread are most favorable on south and southwest aspects because they receive more sunlight and have higher fuel temperatures and lower relative humidity than on north and east-facing slopes. Daytime upslope windsare also stronger on south and west facing slopes (see Slope Winds).

Land configuration

Configuration of land influences the way a fire spreads through slope winds or coastal breezes, eddies at the top of ridges or other sharp topographic breaks, steep canyons, and barriers. As the sun heats land surfaces during the day, air begins to rise on slopes, or convectionally above flat land. Air from cooler locations (lower slopes or the ocean) pushes in behind the rising air, creating upslope winds in the hills or mountains or coastal breezes that flow onshore during the day. The upslope winds and coastal breezes may be major influences on the direction and rate of fire spread during the day. At night, the general air flow reverses to downslope or offshore directions, which can reverse the direction of fire spread.

In mountainous topography, the upslope winds on opposite slopes can slow the spread of a fire as it approaches ridge tops. However, air current eddies over ridges, through passes, and around other prominent land surface features can also spread fire brands and create erratic fire behavior (see Effects of mountains on winds).

In ravines or narrow canyons, fire burning on one side can radiate significant heat towards the other side, drying, preheating and increasing the availability of fuel on the other side (Pyne et al. 1996). Steep canyons may also produce a chimney effect in which a strong convection current is created from unstable air conditions at the surface, drawing air in at the base of a canyon and pushing it out the top. The rapid preheating of fuels and strong convectional winds in narrow canyons may lead to very rapid fire spread from the bottom to the top of the canyon. At times, the entire slope can ignite within minutes (see Valley Winds).

Fires can be controlled or slowed by land barriers. Rocks, bare soil, lakes, streams, roads and trails are all possible barriers.

Several physical and chemical fuel characteristics directly or indirectly affect fire intensity and rate of spread.

Fuel moisture content

Fuel moisture content is important in determining both fire intensity and rates of spread since it controls fuel availability. The more water a fuel contains, the more heat required to ignite and burn. When the moisture level is high in fuel, it may take a long time for the fuel to ignite and spread rates and energy release are low. When the moisture content is low, fuels easily ignite and burn more quickly. Prior to ignition, fuel must be heated to a minimum temperature termed the low heat of combustion. The lower the initial fuel temperature, the more energy is consumed by the fuel in order to raise its temperature enough to initiate volatilization and pyrolysis reactions (see also: Pre-Ignition Phase). This reaction is greatly affected by fuel moisture content. Heat entering the fuel raises the water temperature to 100°C, separates water from the fuel, and vaporizes the water (Whelan 1995). Until all water is vaporized temperatures cant increase further toward ignition temperatures, at which point the fuel begins to release energy in the form of heat.

High temperature, low relative humidity, and low precipitation are ideal conditions for drying of fuel. To see an example of how heat release and fire intensity vary under different fuel moisture regimes, see Heat release rates under different fuel and weather conditions.

Dead fuel moisture is a function of weather conditions and depends on the local environment. As the temperature and humidity change through the day or season, the dead fuel moisture varies. Live fuel moisture depends on the season and correlates with spring flushes and fall curing. Live fuels can serve as a heat sink to slow the spread of fire. However, in times of drought, live fuel that would otherwise not burn becomes a major available fuel source (Reinhardt et al. 1991).

Fuel arrangement

Fuel arrangement is an important factor in both fire intensity and rate of spread. If fuels are arranged too densely, there will not be enough oxygen to maintain combustion, no matter what the fuel size. The higher the packing ratio (tightness of fuels), the lower the combustion efficiency and heat release, and the less fuel is consumed (DeBano et al. 1998).

The horizontal arrangement and distribution of fuels may be continuous, patchy, or broken up by barriers, with significant impacts on the rate of spread. If the fuel exists in clumps, it may be difficult for the fire to spread without strong winds or spotting. Horizontal continuity is especially important for the spread of crown fires (Pyne et al. 1996).

Vertical arrangement describes how fuels are distributed between surface fuels and treetops. A fuel ladder (of shrubs, vines covered with pine needles, and pine crowns) can spread fire from the ground up to the canopy. Old pastures and grass prairies in the southeast are examples of a single fuel stratum; in the Lower Coastal Plain, the palmetto-gallberry-southern pine fuels are at least three strata – a forest floor mat of leaves, debris and pine needles in different stages of decay along with other living forbs, the shrub layer, and the pine canopy (Johnson and Miyanishi 2001).

Fuel size and shape

Fuel size and shape also affects combustion, heat release, and rates of spread. Fine fuels are important for the initial ignition and spread of fires. Due to their large surface area-to-volume ratio, fine fuels ignite more easily and release heat more quickly than large diameter fuels (DeBano et al. 1998). Small fuels are also needed to ignite larger fuels, which are potentially significant heat sources. In contrast, large fuel consumption requires fire to be “in residence” longer in order to penetrate the surface. Because of this requirement, large fuels are primarily consumed by smolderingglowing combustion after the passage of the active flaming front and are less important in determining rates of spread. The density of the large fuels can also affect heat release rates. For example, a rotten log ignites more easily and releases heat at a faster rate than a sound log (DeBano et al. 1998).

Fuel load

Available fuel load (weight of fuel/unit area) is one of the most important factors controlling fire intensity. Fire intensity is directly proportional to a fuels heat of combustion, the amount of fuel consumed, and a fires rate of spread. The total energy released in the fire is proportional to the amount of carbon stored in the fuel that is consumed. This relationship is represented in the fireline intensity equation, where the weight of fuel consumed per unit area in the active flaming zone (W) is directly related to rate of heat release.

Fuel loads are dependent on forest type, life stage of the forest (older, over-mature forests may have an accumulation of large woody debris), and time since last fire (DeBano et al. 1998). The proportion of this total fuel load that is consumed is influenced by fuel availability, which in turn is determined by moisture content, chemical characteristics, and size. Fire type also determines the amount of fuel consumed: a surface fire that consumes only needles and grasses consumes much less fuel than a stand-replacing crown fire. To see an example of how heat release and fire intensity vary under different fuel types, see Heat release rates under different fuel and weather conditions.

Fuel chemical composition

The chemical composition of fuels can affect their flammability, influencing both fire spread and energy output. For example, high mineral element concentrations in woody tissue and leaves reduce the flammability of these fuels, reducing total energy released (DeBano et al. 1998; see Particle Flammability). High concentrations of phosphorus in fuels have been shown to reduce a fire’s rate of spread (Lindenmuth and Davis 1973). Phosphorus reduces flammability to such a great extent that di-ammonium phosphate has been used as a fire retardant in fire fighting operations (Foster 1976).

In contrast to mineral content, fuels with high concentrations of oils, resins, or other volatiles can greatly increase their flammability and heat output due to their high energy content (Whelan 1995). For example, the heat of combustion for oak wood is 19.33 MJ/kg whereas the heat of combustion for pine pitch is 35.13 MJ/kg, an increase of over 80% (McArthur and Cheney 1972). See: Plant flammability.

Fuel type or model

Fuel types and wind effects on rate of spread

Fuel types vary in their fuel moisture content, physical properties, and chemical properties. Therefore rates of spread will also vary according to fuel type, often defined by fuel models.

Fuel type effects on rate of spread are illustrated in the figure for three southern fuel models: open pine with grass understory (FM 2), Coastal Plain shrub communities under pine (FM 7), and logging slash (FM 11). Spread rates were modeled using Behave Plus 2.0, a fire behavior prediction system based on Rothermel’s equation. Total dead fuel loads in the three models are 3.5, 4.5 and 11.5 tons per acre, respectively, with FM 2 having twice as much 1-hour fuels (dead grass and pine needles) as FM 7. The higher fuel load of fine fuels in FM 2 leads to the much higher spread rates in that fuel type, especially as winds increase.

To see an example of how heat release and fire intensity vary under different fuel types, see Heat Release Rates under Different Fuel and Weather Conditions.

• Ground fires: restricted to the layer of duff, roots, and buried or partially buried dead and decaying logs.

• Surface fires: propagate through fuels less than two meters high, which are commonly small trees, shrubs, herbaceous vegetation, and litter.

• Crown fires: advance through the aerial strata of the forest more or less independently of surface fires

• Spotting: occurs when firebrands are transported naturally by wind, convection, or gravity beyond the main perimeter of the fire.

Ground fires are restricted to the layer of duff, roots, and buried or partially buried dead and decaying logs and burn very slowly with little if any flame (see Ground Fuels). As compared to surfacecrown fires, there has been little research on rate of spread in ground fires. Fire spread through ground fire is usually slow because of the compactness of ground fuels, with burning by smoldering combustion. Ground fires in dry organic matter in swamps will often burn out much larger areas beneath the surface than is visible from above ground, creating potentially hazardous situations for people walking through the areas after a fire. Due to these hazards, fire practitioners are often wary of burning wetlands with organic soils. It is important to recognize, however, that ground fires are a natural and necessary disturbance in the maintenance of certain wetland communities in the Southeast (see: Fire Regimes of Pocosins and Large Shrub Bogs and Fire Regimes in the Okeefenokee Swamp).

Surface fires propagate through fuels less than two meters high, which are commonly small trees, shrubs, herbaceous vegetation, fallen leaves, pine needles, and woody litter (see Surface Fuels; Pastor et al. 2003). Surface fires are by far the most common type of fire in the fire regimes of most southeastern communities (see Fire Ecology).

In surface and crown fires, the head is the most rapidly spreading portion of a fire’s perimeter and is usually on the upslope or downwind side of the fire. Fire intensities are higher in the head of a fire than elsewhere around the fire perimeter. Key factors that influence head formation are rate of spread and intensity, which in turn are influenced by wind speed, slope, and fuel type.

Crown fires are defined as fires that advance through canopy fuels more or less independently of surface fires (DeBano et al. 1998). Crown fires are extremely dangerous, very difficult to fight and often require the use of indirect suppression methods (Fieldhouse and Dickinson 2003; Pastor et al. 2003).

Van Wagner (1977) grouped crown fires into three broad classes: dependent, active, and independent based on the degree of dependence on the surface fire. Dependent crown fires are initiated and maintained by the heat produced by the consumption of surface fuels and may consume individual tree crowns or small clumps of trees. They do not spread from crown to crown except for adjacent trees in a clump. An active crown fire runs through both the surface and aerial strata at the same time, while an independent crown fire burns through crowns independent of the surface fire. Rothermel (1991) further classified fully developed crown fires into two categories: ‘wind driven’ (strong surface winds) or ‘plume-dominated’ (massive convection column development creates strong local winds, firewhirls and firebrand ignitions outside the fire perimeter). The transition from a surface fire to a crown fire is dramatic.

While dependent crown fires are common in the Southeast, active and independent crown fires are uncommon in the region. Particular exceptions are the stand-replacing fires common in sand pine scrub forests in central Florida, in stand-replacing fires in non-indigenous melaleuca forests in south Florida, and limitedly in Table mountain pine forests in the southern Appalachians. General exceptions to this pattern occur in fires with extreme fire behavior (caused by low moisture levels, erratic winds, or high fuel loadings).

Beighley and Bishop (1990) list ideal conditions for the development of crown fires as: dry fuel, low humidity with high temperatures, heavy accumulations of dead and downed litter, conifer regeneration and other ladder fuels, steep slopes, strong winds, unstable atmosphere, and a continuous cover of coniferous trees. In the South, dense shrub understories are often the main surface fuel that contributes to initiating and maintaining crown fires in dense pine stands. Propagation of a crown fire after ignition from below depends on the spread rate in, and mass density of, the foliar fuel in the crown layer (Van Wagner 1977). Therefore, if the crown fire does not spread fast enough it will burn itself out. For example, during extreme weather conditions when the wind speed is high and sustained, a crown fire may continue and spread for many hours, covering large areas and crossing barriers which would impede the spread of a fire under normal conditions (Pyne et al. 1996).

Spotting occurs when firebrands (sparks, embers or larger burning pieces of vegetation) are transported naturally by wind, convection, or gravity beyond the main perimeter of the fire. Showers of burning embers that fall within ¼ to ½ mile from the main fire can ignite large areas almost simultaneously creating a pseudo front (Brown and Davis 1973). This mode of spread is most often associated with large wildland fires (Pastor et al. 2003), but can occur anytime there are dry, windy conditions:

• Short-range spotting from firebrands can occur in any fuel type and are more likely to occur when the relative humidity is below 40% and are almost a certainty once the relative humidity reaches 25% or less (Weir 2003).

• Air temperatures above 60ºF (Bunting and Wright 1974) and wind speeds over 20 mph significantly increase the chance of firebrand ignitions (Wright and Bailey 1982).

Fuel types also influence the development of firebrands. Light fuels such as palmetto fronds and Spanish moss, which are easily picked up by convection may be carried long distances before they drop and ignite fresh fuels.

Long-range spotting indicators are large glowing firebrands, moderate winds and well developed convection columns. Convection columns form above a fire, with the height of the thermally-produced ascending column of gases, smoke, and debris dependent on air stability and heat released by the fire. As the column is pushed beyond the fire front, it may significantly contribute to fire spread by greatly increasing the likelihood that firebrands will drop onto fresh fuels and new fires will ignite. When air movement in the convection column interacts with the fire, fire whirls may form.

The growth of fire follows a similar cycle regardless of fire type, size or intensity. Under perfectly uniform conditions wildland fire will move from ignition through an initiation stage, during which rate-of-spread accelerates to a steady state, explained below. Under real-life conditions the fire’s shape and rate of spread will continually change. Fuel, weather, and topography combine to influence a fires rate of spread and overall behavior. In the case of prescribed fires, ignition patterns can affect fire growth and intensity. Rarely, a fire can continue to increase its rate-of-spread and intensity, resulting in extreme fire behavior.

Initiation

The initiation stage of a fire is the period of time after ignition until a fire reaches a steady rate-of-spread. Wildland fire generally starts from a single ignition point (lightning strike, match, spark). During the initiation phase there is a buildup of heat and the fuel present at the point of origin is consumed. As radiant heat dries the adjacent fuels the fire will begin to expand its perimeter outward. The rate of spread accelerates initially, with the rate dependent on fuel type, load, compactness, continuity, and depth. Weather factors also affect initial rates of spread and the rate at which fire perimeter spread accelerates. In general, the characteristics that lead to the highest rates of spread include continuous, fine fuels, loosely packed, with low fuel moisture contents on windy days in late spring with high temperatures and low relative humidity.

Steady state

If all fuel conditions remain constant, the fire perimeter will expand at an accelerating rate for 10 to 30 minutes (depending on fuel type), at which point the rate of spread reaches a steady state. In a steady state, under uniform conditions, the flaming zone moves at a nearly constant rate with the same intensity, simply increasing the area that has been burned. The time required to reach steady state is independent of weather, while the actual rate of spread is a function of weather and fuel types. The steady state rate of spread is maintained as long as fuel, weather, and topographical conditions remain stable. Stable conditions are unlikely to occur in wildland fires.

Wildfires that originate from a single ignition point (lightning, arson, or accidental) are often referred to as point source fires. Fire will spread out from this point in a circular pattern in an area with uniform fuel distribution, on level ground, with no wind present. In the presence of wind, the fire will begin to spread fastest in the direction of the prevailing winds or in cases of steep topography, upslope. This area will become the front or head of the fire. The lateral and upwind (or downslope) portions of the fire will spread at a slower rate and are referred to as the flanks and rear of the fire, respectively. The resulting fire perimeter is an elongated or elliptical shape. The more uniform the fuel, terrain, and weather conditions, the more the fire will resemble an ellipse. However, conditions are usually variable and as the fire perimeter spreads its shape will continually change.

Fuel types, topography, weather conditions, barriers, and spotting can influence and alter the shape of the fire. As spot fires ignite outside the perimeter of the main fire, they are incorporated into the expanding main fire line resulting in ‘fingers’ around the edge of the fire. The effect of spotting on fire growth will depend on the fire intensity. Under extreme fire conditions, firebrands can ignite new fires one or more miles downwind from the original fire perimeter. The distance of the main fire from the spot fire will affect fire shape as the spot fire will likely be drawn into the main fire perimeter by the convective column.

The fire perimeter in fine fuels, such as grasses and leaf litter, will change shape more rapidly and with more variance than in heavier fuels. As wind and slopes increase the length-to-width ratio of the elliptical perimeter will also increase. The uniform perimeter disappears as wind direction and fuel type or load change. Changes in these conditions can cause the fire perimeter to develop U-shaped indentions known as pockets or protrusions of fire known as fingers. Former flanks can become the head or rear of a fire for some period of time. Gradually, the fire perimeter will become very irregular.

With prescribed fires, the pattern of fire spread may be very different because of ignition patterns that use lines of fire or multiple spots. A line ignition is created by continuously dripping fuel from a hand held, ATV-mounted, or helicopter drip torch. The resulting fire includes a line of fire moving with the wind and another line backing into the wind (if not ignited along a fire break). Multiple spot or grid ignitions are a method of igniting fires in which ignition points are set individually at predetermined spacing with predetermined timing throughout the area to be burned. These spot fires will expand elliptically until they burn into each other, with a brief increase in fire intensity as the individual fire perimeters merge. See also: Firing Techniques.

Rate of spread is the horizontal distance that the flame zone moves per unit of time (feet per minute) and usually refers to the head fire segment of the fire perimeter. However, rate of spread can be measured from any point on the fire perimeter in a direction that is perpendicular to the perimeter. Because rate of spread can vary significantly over the area of the fire, it is generally taken to be an average value over some given period of time. The fastest rate of spread is along the forward moving perimeter located at the head of the fire. The slowest rate of spread will be found on the windward (back) side of the perimeter. The rate of spread along the flanks will be intermediate between the head and backing rates of spread (see: Fire shape). Rates of spread can easily be estimated by timing the passage of the flaming front between two landmarks of known distance apart. To determine rate of spread within the interior of a fire, firecrackers placed at known intervals along a transect perpendicular to the flame front have also been used. More technical techniques of measuring rates of spread involve videography or the use of thermocouples to record the passage of the flaming front.

Rate of spread is directly related to the amount of heat received by the fuels ahead of the flaming zone, and the heat is a function of the energy release rate per unit area of fire front (I_R reaction intensity = Btu/ft²/minute, see Rothermels equation below). The rate of spread is a key component in planning and decisions for conducting prescribed fires and suppressing wildfires.

As Rothermel’s equation below demonstrates, rate of spread is strongly influenced by fuels, winds, and topography. Rates of spread generally increase with increasing wind speed, slope, and amount of fine fuels. More information is provided in the following pages:

• Effects of weather on fire intensity and rate of spread

• Effects of fuel on fire intensity and rate of spread

• Effects of topography on fire intensity and rate of spread

When fire growth leads to extreme fire behavior the fire itself becomes a fourth factor that influences rate of spread. Fire can produce enough heat to modify local winds, contribute to atmospheric instability, and cause cloud development (Pyne et al. 1996).

<equation> Rate of spread as a function of energy release was mathematically expressed by Rothermel’s equation which has been the basis for most computerized fire spread prediction models such as Behave Plus. The rate of spread equation is simply the heat received by fuels ahead of the fire (numerator) divided by the heat required to ignite the fuels (denominator):

R=I_R ξ(1+ Φ_W + Φ_S) / ρ_η ε Q_ig

where:

• R = rate of spread of the flaming front
• I_R = reaction intensity (defined in Fire Intensity)
• ξ = proportion of the reaction intensity that heats adjacent fuel particles to ignition
• Φ_W = dimensionless multiplier accounting for the effect of wind in increasing the proportion of heat that reaches adjacent fuels
• Φ_S = dimensionless multiplier accounting for the effect of slope in increasing the proportion of heat that reaches adjacent fuels
• ρ_η = ovendry fuel per cubic foot of fuel bed (lb/ft)
• ε = dimensionless number accounting for the proportion of a fuel particle that is heated to ignition temperature at the time flaming combustion starts (near unity for fine fuels and decreases toward zero as fuel size increases)
• Q_ig = heat of pre-ignition, or the amount of heat required to ignite one pound of fuel (Btu/lb) (Pyne et al. 1996).

Extreme fire behavior is a level of fire behavior that goes beyond human methods of fire control and prediction. The erratic nature of extreme fire behavior presents too serious a risk to the lives of ground crews to perform direct attack. Extreme fire behavior is unpredictable even by complex fire behavior models (Albini 1976, Hough and Albini 1978, Rothermel 1983). Compared to western regions of the U.S., extreme fire behavior in the Southeast is uncommon and generally only occurs during periods of severe drought. Under these conditions, communities where extreme fire behavior is most likely to occur include pocosins, palmetto-gallberry flatwoods, Table Mountain pine, and sand pine scrub.

Although extreme fire situations are few in the Southeastern U.S., they account for most of the fire-related injuries and fatalities. Below, the characteristics of large/mass fires are described and the fuel, weather, and topographical factors that create extreme fire behavior are explained. These concepts are then combined in two cases studies of extreme fire behavior that occurred in the Southeast: the Air Force Bombing Range Fire of 1971 and the Florida wildfires of 1998.

Characteristics of extreme fire behavior include:

• very high to extreme rates of spread;
• prolific crowning and torching;
• fire whirls;
• tall, well-developed convection columns;
• long flame lengths; and
• excessive spotting (National Wildfire Coordinating Group 1981).

High rates of spread and long flame lengths are the most common types of extreme fire behavior. Rates of spread are dramatically altered by long distance spotting. A fire may begin exhibiting the characteristics of extreme fire behavior during an event known as a blowup, which is a sudden increase in the intensity and rate of spread of a fire. Blowup occurs when windspeed increases dramatically and/or an area of atmospheric instability settles over the fire. Blow-ups are difficult to control and have taken the lives of wildland firefighters. They generally subside with a change in fuel or weather conditions.

One of the more fascinating manifestations of extreme fire behavior is fire whirls. Fire whirls are spinning vertical vortices of hot air and gases rising up from a fire that carry smoke, embers, debris, and flame. They can range in size from less than one foot to more than 500 feet in diameter. Large fire whirls have the intensity of a small tornado and can rapidly spread embers and new fire spots beyond the perimeter of the main fire.

Torching and crowning are also spectacular examples of extreme fire behavior. Torching occurs when an entire tree, or clump of trees, goes up in flames, but the fire does not spread to the adjacent crowns. In contrast, a crown fire occurs when fire spreads from crown to crown in the overstory. Potential for torching and crown fire initiation increases as the height to the base of the crown decreases or ladder fuels increase. Ladder fuels (vines, small trees, large shrubs) provide vertical continuity between surface fuels and tree canopy fuels.

Excessive spotting is also indicative of extreme fire behavior and can be linked to strong winds, the presence of fire whirls or a convection column. Spotting occurs when sparks, embers, and/or firebrands are carried up into the convection column and dropped outside of the perimeter of the original fire, causing an entirely new fire. Embers are small pieces of burning material that usually go out before they spread more than ¼ mile from the ignition zone (National Wildfire Coordinating Group 1981). Firebrands are larger materials like branches or palm fronds, which can be carried much further while still burning and may cause new fires more than ¼ mile from the original fire (National Wildfire Coordinating Group 1981). In 1983, Albini published mathematical equations that predicted the maximum height and drift of firebrands. These equations were simplified for easy field calculations and were further refined in a publication by Chase in 1984.

Extreme fire behavior results from combinations of fuel and weather conditions that include:

• abundant available fuels,
• drought,
• low relative humidity,
• strong winds,
• atmospheric instability, and
• topography (in mountainous areas).

Fuels

High loading of available fuel affects the ability of wildfire to develop into extreme behavior. The availability of fine dead fuels and live foliage depends largely on their moisture content. Very low fuel moisture content allows wildfires to consume more of the fuel and therefore release more heat. Larger dead fuels will also cure under drought conditions, and can become important heat sources during extreme fire conditions.

Some fuels are more prone to long-range spotting: medium to heavy palmetto-gallberry, dry swamps, and canopies with crown closure greater than 75%. Certain fuel types also increase the incidence of extreme fire behavior. For example, the presence of foliage with a profusion of volatile oils such as gallberry (Ilex glabra) can increase the incidence of torching and crowning into the canopy.

Weather factors

Extreme fire behavior is likely to ensue if ignition occurs when fine fuel moisture content is below 5% and the relative humidity is below 20% (National Wildlife Coordinating Group 1989). Wind speeds above 20 mph can produce extraordinary rates of spread under droughty or dry conditions.

Atmospheric instability can play a role in the evolution of extreme fire behavior. Atmospheric instability contributes to the development of a convection column by allowing rapid vertical lift in the rising column of hot gases and debris (see Fire and Plumes). The influx of air at the base of the convection column causes erratic winds and rapid fire growth. The Haines Index is used to measure the degree of atmospheric instability and dryness of the air over a fire. An index of 6 indicates that there is high potential for the type of fire growth that is observed in blow-ups. Another indicator in the South is the dispersion index which measures the process by which smoke mixes with the atmosphere and is carried away. A daytime dispersion index greater than 60 is an indication that wildfires will be difficult to control; values less than 30 will present smoke and visibility issues, but not erratic fire behavior. Dispersion index values in excess of 100 were reported during the 1998 fires in Florida.

Drought conditionshave a role in promoting extreme fire behavior by drying out and curing fuels. The Keetch-Byram drought index (KBDI) estimates the dryness and thus the flammability of the duff and soil horizons (Pyne et al. 1996). It takes into account the net effect over time of precipitation and evapotranspiration on the amount of moisture in these horizons. A value of 0 indicates that the duff and soil horizons are completely saturated with water while a value of 800 indicates that it would take eight inches of rain to lower the KBDI to zero (or saturate the soil). The index increases every day without rain and decreases when it rains. The amount it increases and decreases depends on the types of soils found in a region. Indirectly, the KBDI is also related to the gradual drying of large dead fuels, which contribute significantly to extreme fire behavior.

Air Force Bombing Range Fire

A well-documented wildfire exhibiting extreme fire behavior in the South was the Air Force Bombing Range Fire which occurred in eastern North Carolina on March 22, 1971 (Wade and Ward 1973). The fire crowned through 15,000 acres of pond pine with pocosin understory during the first 4 hours and a total of 29,300 acres in four days. It had an average rate of spread of 2 mph and a peak of 5 mph. Low relative humidity for four of the five days preceding the start of the fire cured the fuels. A second long run occurred the first night of the fire during the passage of a dry cold front, with wind gusts up to 40 mph and low relative humidity. The fire died down only when its flaming fronts ran into wet, marshy areas (Wade and Ward 1973).

1998 Florida Fires

Florida was affected by an unusually strong ENSO (El Nino-Southern Oscillation) in the year prior to the 1998 wildfires, which helps to account for their severity. ENSO is composed of three phases: El Nino, La Nina, and neutral. In Florida, the El Nino phase produces higher than normal levels of rainfall while the La Nina phase produces lower than normal rainfall. ENSO was in an El Nino phase during the winter of 1997. As a result of the increased rainfall, there was an abundance of new plant growth in many ecosystems. However, these plants dried up in the spring of 1998 when ENSO switched suddenly to a La Nina phase, creating a severe drought. During the 1998 Florida wildfires, the Keetch-Byram drought index (KBDI) reached an all-time high of 780. The curing of the new plant growth contributed to the already hazardous fuel loads found in many forests in 1998. Fuel loads over 10 tons/acre are considered hazardous (Wade 1989). With high fuel loads and droughty conditions, the state was susceptible to both arson and summer lightning storms. Approximately 500,000 acres in Florida burned in more than 2000 wildfires in May and June 1998 (Butty et al. 2001).

Fire whirls are a rare but potentially catastrophic form of fire. These buoyant fire plumes are columns of ascending air with rotational motion which carry aloft smoke, debris, and flames greatly increasing the rate of fire spread (Pyne et al. 1996). Fire whirls range from a foot or two in diameter to small tornadoes in size and intensity (Brown and Davis 1973). Once a fire whirl leaves the influence of the main fire it quickly loses intensity and becomes an ordinary whirlwind (Schroeder and Buck 1970).

In order for a fire whirl to exist, there must be an organized source of angular momentum, either from wind shear or from the fire’s convection column, which in turn creates large swirl velocities as air is entrained into the fire plume (Battaglia et al. 2000). Firewhirls occur most frequently where heavy concentrations of fuels are burning and a large amount of heat is being generated in a small area. Mechanical forces are often present which serve as triggering mechanisms to start the whirl (Schroeder and Buck 1970). A common area for firewhirls is the lee side of a ridge where the heated air from the fire is sheltered from general winds. Mechanical eddies produced as the wind blows across the ridge can serve as the triggering mechanism to initiate the firewhirl. On flat terrain they tend to occur on the lee side of the fire near the outside edges of the front (see Vortex Phenomena).

Even in small fires that arent exhibiting extreme fire behavior, fire whirls can be created by lighting patterns. A rotational motion can be generated when two lines of fire come together, particularly if they approach at right angles. If large firewhirls are created, fire control problems can result.

A vortex is a convective eddy that spins about either a horizontal or vertical axis. The air moves around a central axis as well as toward and along the axis. Vortices occur at a variety of scales in fires, varying from only a few inches or feet across that last a few seconds to several hundred feet in diameter and thousands of feet high. They play a significant role in movement of the fire front and in spreading fire beyond the fire’s perimeter, especially in large fires. For example, horizontal roll vortices are a major mechanism in crown-fire spread, and post-burn aerial photography indicates that unburned tree-crown streets are common with crown fire (Haines 1982).

Vortex formation often begins with horizontal rolls caused by wind speeds changing with height (vertical shear). When the horizontal vortex crosses a local region of buoyancy (as in a convectional updraft), a portion of the vortex is tilted up, and often split, creating two vertical eddies spinning in opposite directions. Vortex formation can also occur without outside forces such as wind shear if gradients of pressure or density occur within a plume in opposing directions. A vertical vortex intensifies as air moving into the vertical “pipe” is lifted up (or forced down), causing the vortex to stretch and narrow, further increasing the speed of air parcels within the column (Jenkins et al. 2001).

There appear to be two types of fire whirlwinds in the development of large vortices: one type occurs as a pair on the leeward side of the convection column and the other type is created after the entire convection column begins to rotate. The second type occurs in association with very intense fire that may be described as fire storms.

In normal fire conditions, smoke rises and disperses without noticeable impacts on fire behavior. Rates of spread are governed by radiation and convection heat transfer, air movement in and above the fuel bed, and characteristics of the fuel bed. If fires begin to burn very intensely, air movement in and around the smoke column begins to impact the fire itself, resulting in plume-dominatedconvection-dominated fires.

Plume- or convection-dominated fires are characterized by high rates of heat release, and unpredictable and extreme fire behavior. These types of fires create dangerous conditions for fire fighters. The formation of plume-dominated fires is correlated with a high value on the Haines Index. How convection columns and plumes develop and how they can dominate fire behavior is discussed below.

Convection and plume development

A plume is characterized by rising buoyant gases, with the buoyant forces provided by intense localized sources of heat. Plume rise is a function of free convection in the atmosphere. As a fire heats and expands air near the ground, large density differences between the heated volume and the surrounding air mass are created, causing the heated parcel to rise. The potential height of the resulting plume depends on the heat energy of the source and rise velocity, which is affected by air stability, the exchange and conservation of mass, radiant heat loss, the buoyancy force, and turbulent mixing with the ambient air.

Atmospheric stability affects the size, shape, and growth of smoke columns. Stable air generally discourages height growth of smoke plumes because the rising heated air cools faster than the warm ambient air and temperatures tend to equalize at a relatively low “mixing height” in the atmosphere. Without a strong vertical movement of air, fire activity is generally lower with stable air masses unless there are strong surface winds. Unstable air masses create conditions for rapid vertical plume development and convectional lift, which increases surface air movement toward the fire. The convection column created by these conditions contains ascending columns of hot gases, smoke, and entrained debris.

Smoldering fires often create plumes that are neutrally buoyant, limiting widespread dispersion but allowing surface winds to dominate smoke trajectories. This can lead to accumulations of smoke in valleys and basins at night. Inversion layers (including marine or subsidence inversions) act to inhibit vertical motion in the atmosphere and affect smoke dispersal (Pyne et al. 1996; Potter 2002). Smoke will not dissipate in these conditions and can cause thick fog-like conditions that hamper visibility.

A fire is said to create its own wind when the convection column intensifies to a level such that air rushing into the column base to replace the air leaving the top of the column is stronger than the ambient wind field. Plume rise calculations are essential for determining smoke dispersion patterns and possible effects on fire behavior. Flame heights may increase in such convection columns and ascending gases from volatile materials may be completely consumed. Increases in combustion and convection act as feedback mechanisms on each other and can lead to convection columns developing to heights of 25,000 to 50,000 feet. The rush of air entering a convection column may transport firebrands up and out of the column causing spot fires to break out under severe conditions. Despite the volume and height of the smoke, smoke quantity near the base of the column may remain minimal particularly on level ground. In areas where topographic gradients exist, the slope may cause the convection column to tilt, creating low visibility upslope and increasing the rate of spread as firebrands fall out of the column.

High intensity fires can develop central convective columns with counter-rotating vortices that involve massive entrainment of the surrounding air mass (Clark et al. 1996; Haines and Smith 1987; Haines and Updike 1971). This stage of fire can produce fast-rising plumes and turbulent downdrafts that can exacerbate an already dangerous situation. The downdrafts and fallout from the column carry sparks and embers that ignite new fires. Cumulonimbus clouds often develop with accompanying lightning and rain. Dynamic plume rise brings gas and particles high into the atmosphere where strong winds can disperse the smoke hundreds to thousands of kilometers. As high intensity fires cool, however, the central column often collapses, creating numerous small convective cells that are less dynamic but equally active in carrying smoke into the atmosphere.

Diverse smoke management models are currently in use by various private and governmental agencies. These models were developed from basic algorithms developed in the mid-1960s to describe plume rise for industrial ducted emissions (see Background of plume models). Current models include such data as temperature of the fire, land area, flaming regions, fire plume duration, and emissions load (see Harms and Lavdas 1997, Jenkins et al. 2001, Latham 1993, Linn et al. 2002, Mercer and Weber 2001). Some recent models incorporate weather parameters such as wind speed, vertical lapse rates, barometric pressure, and moisture levels at varying altitudes to determine the extent to which atmospheric variation will affect smoke dispersion.

Plume Models

One of the simplest ways of estimating smoke concentrations is to assume that plumes diffuse in a Gaussian pattern along the centerline of a steady wind trajectory. Plume models usually assume steady-state conditions during the life of the plume, which means relatively constant emission rates, wind speed, and wind direction. For this reason, they can be used only to estimate concentrations relatively near the source or for a short duration. Their steady-state approximation also restricts plume models to conditions that do not include the influence of topography or significant changes in land use, such as flow from a forest to grassland or across a land-water boundary.

Plume models typically are in Lagrangian coordinates that follow particles or parcels as they move, assigning the positions in space of a particle or parcel at some arbitrarily selected moment. Examples adapted for wildland biomass smoke include VSMOKE (Harms and Lavdas 1997; Lavdas 1996) and SASEM (Riebau et al. 1988; Sestak and Riebau 1988). VSMOKE assumes that wind and other weather conditions are steady and constant over the area of consideration and during the time smoke will move from source to receptor. To predict smoke movement and dispersion, VSMOKE assumes that mean wind flow has dispersion capabilities due to wind movement caused by small random right angle deviations from the mean flow. As an example, the mean wind may be blowing from the west, with deviations from the south and north (causing horizontal dispersion). Additionally, right angle movement of the wind may be from below or above causing vertical dispersion.

Puff Models

Instead of describing smoke concentrations as a steadily growing plume, puff models characterize the source as individual puffs being released over time. Each puff expands in space in response to the turbulent atmosphere, which usually is approximated as a Gaussian dispersion pattern. Puffs move through the atmosphere according to the trajectory of their center position. Because puffs grow and move independently of each other, tortuous plume patterns in response to changing winds, varying topography, or alternating source strengths can be simulated with some accuracy. Most puff models are computed in Lagrangian coordinates.

Particle Models

In a particle model, the source is simulated by the release of many particles over the duration of the burn. The trajectory of each particle is determined as well as a random component that mimics the effect of atmospheric turbulence. This allows a cluster of particles to expand in space according to the patterns of atmospheric turbulence rather than following a parameterized spatial distribution pattern, such as common Gaussian approximations. Therefore, particle models tend to be the most accurate way of simulating concentrations at any point in time. Particle models use Lagrangian coordinates for accurate depiction of place for each period of particle movement (for example, Hysplit: Draxler and Hess 1998; PB-Piedmont: Achtemeier 1994, 2000).

Grid Models

Grid models use Eulerian coordinates, disperse pollutants uniformly within a cell, and transport them to adjacent cells. The simplicity of advection and diffusion in a grid model allows these models to more accurately simulate other characteristics of the plume, such as complex chemical or thermal interactions, and to be used over large domains with multiple sources. This is why grid models commonly are used for estimating regional haze and ozone and are often called Eulerian photochemical models. Much of the future work on fire impact assessment and planning at regional to national scales will be done using grid models.

Because of their nature, grid models are not used to define accurate timing or locations of smoke concentrations from individual plumes, only concentrations that fill each cell. This means that sources small relative to the grid size, which create individual plumes, will introduce unrealistic concentrations in places that are outside of the actual plume.

The basic mechanisms and algorithms used to describe plume rise and buoyancy were developed in the mid-1960s by Briggs (1969) for industrial, ducted emissions. These methods are still used today to estimate the plume rise and buoyancy of smoke from fires in spite of the significant differences in characteristics between ducted emissions and prescribed and wildland fires:

• Heat released from ducted sources is precisely known and usually emitted at relatively constant rates during a single phase of combustion. Heat released from fires is a function of fuel loading, fuel conditions, and ignition method through several phases of combustion (pre-ignition, flaming, smoldering, and residual), which create highly variable magnitudes and rates of heat release.

• Nearly all of the energy generated at the source of a ducted plume is transmitted to convection energy. In open burning, however, significant amounts of energy are lost by conduction and radiation, reducing the amount of available energy for convection.

• Plumes from ducted sources create single convective columns, but low intensity understory burning that occurs over broad areas does not develop a cohesive plume.

To improve plume rise predictions, emission production models need to characterize the spatial and temporal pattern of heat release from fires, and plume rise models need to account for the energy lost from the convective system through radiation and turbulent mixing. Models such as EPM and Burnup simulate variable rates of heat release from fires, but they use general estimates of spatial distributions of fuel, including structure, composition, and moisture content. Also, significant elements of fires that influence convective energy—such as the distribution of naturally piled fuel (“jackpots”), amount and density of rotten fuel and duff, and release of water vapor—are not adequately captured.

Low intensity fires that typically do not have a cohesive convective column have also been difficult to model. These fires have been treated, from a modeling perspective, as an area source in Eulerian grid models. However, since these models use Eulerian coordinates that are fixed in space and time, there is no attempt to identify individual particles or parcels from one time to the next. In Lagrangian dispersion models, there is currently no valid means of calculating plume rise from unconsolidated convection. Lagrangian models (bell-shape or Gaussian distribution pattern, often applied to plume and puff models) are used to show smoke concentrations crosswind of the plume.

Another complication for modeling is that once plumes from fires enter the atmosphere, their fluctuating convection dynamics make them more susceptible to erratic behavior than well-mannered industrial stacks. For example, different parts of a plume can be carried to different heights in the atmosphere at the same time. This causes unusual splitting patterns if there is a notable wind shear between elevations, causing different portions of the plume to be transported in different directions. Therefore, predictions of the plume’s impact on visibility and air quality under these conditions become highly uncertain (Walcek 2002). Even when the behavior of plumes from fires resembles that of stack plumes, the varying and widely distributed locations of wildland sources prevent consistent study. For example, down-wash of plumes has been observed from ducted (stack) emissions after an inversion breaks up — conditions that are common at the end of an onshore breeze if the plume is above the inversion at its source (de Nevers 2000; Venkatram 1988) or if horizontal stratification in the lower atmosphere is disrupted by mountains (de Nevers 2000).

The Haines Index is a national fire-weather index based on the stability and moisture content of the lower atmosphere and their direct relationship to the growth of large fires (see Extreme Fire Behavior and Fire and Plumes). The index ranges from 2 to 6, with 6 being the highest potential for large fire growth (Pyne et al. 1996). The more dry and unstable the atmosphere the higher the Haines index value. The Wildand Fire Assessment System produces daily maps of the Haines Index for the U.S.

Several studies have noted deficiencies in the Haines index (Jenkins 2002; Porterie 2001; Potter 2002). It does not incorporate wind variables, and it may only use weather data from widely distributed locations. Environmental conditions could exist whereby a low Haines value would be computed but local wind direction and speed could create a situation conducive to flare-ups and spotting. In addition, upper level atmospheric moisture is not accounted for with the Haines index. Moisture levels above 10,000 feet do not reflect relative humidity and may increase instability in the upper atmosphere leading to thunderstorm development. Jenkins (2002) suggested the Haines Index scale should be expanded past the value of 6 for the best predictive models of fire behavior as influenced by atmospheric conditions.

Wildland fire displays some remarkable characteristics and behavior. Flames are one of the visible manifestations of the combustion process and are the main source of heat from a fire. Measuring or estimating flame height, length, depth, angle, and char height can help to understand fire intensity and predict fire effects (Johnson and Miyanishi 2001).

Two theoretical definitions of flame height are: the height where flames reach the maximum temperature, which is difficult to measure in wildland fires because of varying flame heights; and the vertical distance from the flame base or ground to the time-averaged yellow flame tip as shown in the figure (Johnson and Miyanishi 2001).

Flame characteristics

Measurements of the visible flame tip require height markers positioned in the fire and either ocular observation or videotaping a moving fire (Adkins 1995, Hamins et al. 1996). A motor-driven 35mm still camera or a video camera can be used to record the fire for measuring characteristics such as flame angle and height (Johnson and Miyanishi 2001). For a smoke-covered flame, an infrared camera can be used to penetrate the smoke and measure the maximum flame temperature and visible flame tip (Johnson and Miyanishi 2001).

Flame length is another estimator of fire intensity. Flame length estimates have been made for many experimental fires, wildfires and prescribed burns. In the South, flame lengths may vary from less than 1 ft in slow moving backfires in leaves and short grass to over 100 ft in extreme wildfire conditions in dense shrubs and pines. Flame length is measured from ground level from the mid-point of the base of the flame to the tip of the flame, whether the flame is vertical or slanted.

Flame characteristics

In theory, flame length has been defined as the distance from the fuel surface to the point on the flame axis where the CO concentration is 5000 ppm. In wildland fires, it is difficult to determine these theoretical parameters and measurements are generally based on video or time-lapse photography. Flame length also provides an estimate of the difficulty in fighting fires. For instance, flame lengths less than four feet can be attacked at the head of the fire, while fires with flame lengths over eleven feet should not be controlled from the head and often lead to crowning and long range spotting.

Flame depth

Flame depth is the distance from the front to the rear of the flame zone at the fuel surface. Flame depth is dependent on the spread rate of fire, the length of time flames remain over any particular point in the fuel bed (residence time), and the type and quantity of fuel being burned.

Tilt or flame angle

Tilt or flame angle is the angle between the unburned terrain whether flat or sloped and the leading surface of the flame. Flame angle is highly influenced by wind. For a wind-driven flame the angle is less than 90^o and for a flame backing into the wind it is greater than 90^o.

Bark char height is the height from ground level to the highest point on the bole where bark char is evident (outer bark blackened by fire). Char height on individual trees is measured halfway between the upper and lower char heights on opposite sides of the tree. Bark char height is used to approximate flame height or flame length, particularly when measuring these variables during the course of a fire is impractical or impossible. Flame length estimated from bark char height can then be used to reconstruct fireline intensity using Byrams equation (e.g., Waldrop and Van Lear 1984). However, evidence suggests bark char height is not a reliable estimate of flame height, flame length, or indirectly, fireline intensity. For example, Cain (1984) found that for both back fires and head fires in a 9-year-old loblolly-shortleaf pine stand, bark char height underestimated flame length by 50 percent and fireline intensity by 75 percent during prescribed burns. Cain noted, however, that bark char height may provide adequate post-burn measures of relative, but not absolute, fire intensity where better measurements are not possible. In fact, Cain found that crown scorch height estimated fire intensity no better than bark char height.

Bark char height has also been use to predict fire damage. For example, Waldrop and Van Lear (1984) found a direct relationship between bark char height and crown scorch above a threshold value of 3 feet in a loblolly pine plantation. However, since bark char is not a very reliable indicator, it should only be used in cases where attempts are made to reconstruct fire behavior months after the scorched foliage has fallen (Wade and Johansen 1986). The accuracy of bark char height is particularly questionable in previously burned stands because sloughing of charred bark sometimes requires several years (Cain 1984).

Computer models can be used to predict fire behavior based on differences in fuels, weather, and topography. These models are used to support fire management decisions, as a training tool to improve fire management skills, and can help display and explain fire behavior and fire management strategies to the general public. While these prediction models have inherent limitations and will never replace the knowledge and experience of wildland managers, they can provide tools to estimate fire behavior and possibly save time, money, and lives. Many of these fire behavior prediction systems require descriptions of fuel properties as inputs for calculations- this information is usually provided by fuel models. Below are examples of both current and developing fire behavior computer models.

• BEHAVE Plus Fire Modeling System. This freeware is the most commonly used software tool for predicting fire behavior in wildland fuels. Originally written in FORTRAN in 1984, the most recent version has user-friendly graphical interfaces that estimate wildland fire behavior under various fuel, weather, and topographic situations. It is used for real-time predictions of behavior of wildfires or prescribed fires. The related fireLib, which uses BEHAVE algorithms is for programmers who need a highly optimized API for developing fire behavior growth simulators.
• FARSITE (Fire Area Simulator) is used to predict fire spread in two dimensions on a landscape (as opposed to BEHAVE, a one-dimensional model). FARSITE requires GIS data for model creation. It is designed for use by those familiar with fuels, weather, topography, wildfire situations and requires training for its use.
• EMBYR is a probabilistic model that predicts potential burn patterns of large fires relative to variations in fuel types and weather patterns in an area. Ignitions can occur at random points or specific locations and ignitions from firebrands can be simulated relative to fuel type. EMBYR requires a GIS layer of fuel types based upon age classes and species composition. Fire spread probabilities are specified for three possible fuel moisture conditions; wet, intermediate, or dry. Probabilities are then adjusted using one of three wind speed categories and one of eight wind directions. The output from EMBYR indicates the final burn pattern of one or more potential landscape-scale fires allowing impacts from future fires to be estimated. The principal drawback to EMBYR is the tremendous diversity in forest species composition and age classes at varying fuel moisture conditions on a large scale thereby making probabilistic map predictions difficult.
• DYNAFIRE requires GIS data layers of fuel types, elevation, slope, aspect, and stream channels. Fuel types are standard fuel models as used in Behave Plus. Elevations can be input from DEMs and slope and aspect are computed from the elevation data. Stream channels can be digitized from topographic maps and this information is related to the elevation layer to indicate fire growth directions. Additional DYNAFIRE inputs include temperature, relative humidity, wind direction and speed, and 10-hour, 100-hour, and live fuel moistures. DYNAFIRE output maps show hourly time contours for fire perimeter, fire line intensity, average fire spread rates, and fire spread direction.
• The LANDFIRE project is one of the most massive undertakings in computer based mapping and modeling. It’s a cooperative effort between the USGS and the US Forest Service. Included in this project are 18 years of US daily weather data, 13 layers of GIS data at 30m resolution, and a host of modeling applications and tools (Keane et al. 2003). A prototype has been funded through April 2005 with final applications due at that time.

• NEXUS is an Excel spreadsheet that links surface and crown fire prediction models. It is used to estimate surface, transition, and crown fire behavior, generate site specific indices of torching and crown fire potential, build and test custom surface fuel models, and evaluate alternative treatments for reducing risk of crown fire.

• FireFamily Plus is a fire climatology and occurrence program that combines and replaces the PCFIRDAT, PCSEASON, FIRES, and CLIMATOLOGY programs into a single package with a graphical user interface for the PC.
• FOFEM (a First Order Fire Effects Model) is a computer program for predicting tree mortality, fuel consumption, smoke production, and soil heating caused by prescribed fire or wildfire. Individual tree species and relative abundance are part of the model input.
• Geo-spatially referenced methods have been explored whose end purpose is the provision of real-time, or very near real-time, mapping of wildfires and associated flame fronts. Satellite imagery has been considered and attempted as a tool for mapping and predicting wildfire spread. While the imagery collected is useful for mapping applications that aren’t time sensitive (Quayle et al. 2003), the lag time involved between overhead passes has proven too great to be effectively integrated in any wildfire attack plan. A more promising application is the Phoenix airborne infra-red fire detection and mapping system. Among its capabilities are infra-red, roll-corrected, geo-spatially referenced photographs that can be sent to the field and downloaded on laptops as jpegs in near real-time. It has the capability of automatically detecting 15-20cm hotspots from 3000 meters altitude traveling at 200 knots. It can cover 40.5 square kilometers per minute and its geo-information can be imported into GIS formats and applications (Smith and Chamberlain 2003).

Computer models are considered valuable tools to assist fire managers with the decision making process. However, quantification of natural processes related to changing weather patterns is difficult at best. Historically, computer generated models have had difficulty simulating several important conditions related to fire behavior. Four of these very influential, but elusive quantifications are:

1. Horizontally discontinuous fuels
2. High frequency variability in winds
3. Interaction of multiple fire fronts
4. Flame attachment to slopes and fire moving through constricted canyons (Finney 2003).
5. Behavior of crown fires
6. Conditions that might lead to fire whirls.

Recent advances in personal computers and GIS technologies are creating opportunities for natural resource agencies to acquire and develop more sophisticated fire behavior models. Still, supercomputers designed to process trillions of calculations per second become bogged-down in the minutiae involved in predicting fire behavior via weather-related variables.

It is understandable that the inherent dynamic nature of atmospheric conditions presents unusual difficulties related to both data collection and input, and the computational downtime associated with new or changing data sets in the fire behavior modeling process. Following are discussions of the inherent difficulties involved with incorporating weather and topographical effects into fire behavior models.

Weather effects in models

The effects of certain weather conditions on wildland fires in ideal situations are fairly well understood. Localized on-site observations of current conditions have served to assist firefighters in predicting short-term fire behavior for many years. Models have incorporated pertinent weather and environmental variables with varying degrees of implementation and utility. Accuracy of forecasts is, however, plagued by the inherent difficulty that arises from quantifying the dynamic nature and relationship of weather conditions to fire behavior. Oftentimes, weather data is either too local for application to a large fire, or too remote for predictions at a localized situation.

It is unrealistic to expect time-sensitive weather predictions to be generally available with such specificity as to address the multiple conditions that may be present at, or near, a wildfire. While local topography and fire/weather relationships have plagued the utility of regional forecasts, the advent of GIS and advancements in fluid dynamics modeling software has shown promise in the realm of scalable weather/topography interactions (Forthofer et al. 2003)

Topography effects in models

Topographical influences incorporated by fire behavior simulation and prediction models have become fairly advanced. The advent of GIS and satellite-based terrain mapping has streamlined the data entry process and provided a high degree of error control. Such in-depth topographical mapping provides reliable probabilities concerning fire spread and direction. Alternatively, the downside of such data intensive methods is that while accurate, they often serve to complicate modeling intentions and diminish the feasibility of such applications in time sensitive settings.

Empirical laboratory testing has allowed operational models to be formulated based on effects generated by a gamut of environmental variables. These models function fairly reliably in low wind, flat terrain situations. Alternatively, physical models have been developed in response to erratic fire behavior that occurs in other than “ideal” conditions. These models base their predictions on physical combustion process formulas designed to mathematically compute the probable fire behavior in a variety of situations. The major drawback to these models is the immense computational downtime. These models are so data driven and relational calculation intensive that supercomputers fail to provide timely results (Achtemeier 2003).

The most commonly used software tool for predicting fire behavior in wildland fuels is the BEHAVE Plus Fire Modeling System (freeware), which predicts fire rate of spread in one-dimension. BEHAVE Plus can be used during fires for real-time predictions of either wildfire or prescribed fire behavior. It can also be used in fire behavior training, prescribed fire planning, and dispatch of initial attack fire crews.

The original BEHAVE system was a series of interactive computer programs for estimating wildland fire behavior under various fuel, weather, and topographic situations. Behave was updated in 1998 (BEHAVE Plus) to reflect changes in both user needs and technology. The updated version uses a more user-friendly windows interface, as compared to the original 1984 version written in FORTRAN. Fuel characterization was expanded beyond the fine fuels that carry surface fire spread to include all components: overstory, organic ground fuel, and surface fuel, including the heavy dead and down fuel that burns behind the fire front. BEHAVE Plus utilizes basically the same fire modules as the old Behave: (1) SURFACE - Surface fire spread, intensity, flame length; (2) SIZE - Area and perimeter of a point source fire; (3) SPOT - Spotting distance; (4) SCORCH - Scorch height; (5) MORTALITY - Tree mortality; (6) IGNITE - Probability of ignition; and (7) RH - Relative humidity determination. Other fire models such as crown fire spread, transition to crown fire, large fuel burnout, smoke production, and soil heating are also now included. Worksheets guide the input of values based on the selected modules, input options, and output variables. Values are input directly or by utilizing input "Guide screens" which display a valid range of values and suitable increments for continuous variables and model choices such as grasslands or brush for discrete variables. Units can be entered in either the English or Metric system.

The weather inputs in BEHAVE Plus include a measure of the wind speed (twenty feet above the ground and averaged over ten minutes, midflame or eye-level wind speed, and a wind adjustment factor which accounts for whether or not fuels are sheltered from the wind) and direction (specified as upslope, degrees clockwise from upslope, or degrees clockwise from North). Additional weather inputs are air, wet bulb, and dew point temperatures. Slope steepness (0-45 degrees) and aspect (0-360 degrees clockwise from North) are entered into BEHAVE Plus in the terrain section of the worksheet. Additional parameters entered into the terrain section include site elevation, ridge-to-valley elevation difference, ridge-to-valley horizontal distance, and spotting source location. The Map section of the worksheet uses input values to determine distances and slope steepness using four variables - (1) map representative fraction; (2) contour interval; (3) map distance; and (4) number of contour intervals.

The BEHAVE Plus version 2.0, released in 2003, added modeling capabilities for safety-zone size, containment, lightning ignitions, dynamic palmetto-gallberry fuel models, and diagrammatical output. It is expected that additional features will be added in future releases.

The modeling capabilities of BEHAVE Plus were used in recent research to illustrate the potential fire behavior in pine flatwood forests after treatment with one of three silvicultural techniques at five different ages-of-rough (Brose and Wade 2002). In each stand, understory fuel characteristics were collected as inputs to BEHAVE Plus to build custom fuel models for each treatment/age-of-rough combination. To match the drought conditions of the 1998 wildfires in Florida, fire behavior was simulated for each treatment/rough-age combination under drought as well as normal weather conditions. Fuel models were combined with weather data to calculate specific fire behavior estimates for each combination (Brose and Wade 2002).

A fuel model is a mathematical representation of one of several natural fuel conditions that is used in fire behavior prediction systems or fire danger rating systems. Specifically, these models describe vegetation in terms that are required by the Rothermel fire spread model (for example, surface area to volume ratio, fuel load, size, shape, compactness, horizontal and vertical continuity, moisture of extinction, etc.).

A series of fuels model classifications systems exist. These systems evolved from a simple conceptual classification of fuels into four groups (grasses, brush, timber, and logging slash), to a 13-model system initiated by Rothermel in 1972 (FBPS fuel models), to a 20-model system developed for the NFDRS in 1978. The most recent advancement in fuel model classification is the Fire Characteristic Classification System (FCCS) system currently being developed by the USDA Forest Service.

• Four basic fuel types: grasses, brush, timber litter, and logging slash.
• FBPS fuel models: Fuel models in this 13-model system, developed by Rothermel in 1972 and later expanded by Albini, are used in fire behavior prediction systems (FBPS) such as BEHAVE.
• NFDRS fuel models: These fuel models, similar to the FBPS fuel models but with extra parameters needed for fire danger prediction, are typically used in fire danger rating systems such as the NFDRS.
• Fire Characteristic Classification System: This system offers a more comprehensive description of a fuel bed than earlier fuel models. A prototype system currently classifies over 200 fuel types in the US.

Grass fuel models range from very fine porous fuels to continuous herbaceous fuels that are cured or nearly cured. Grasses have high surface area to volume ratios, can lose or gain moisture quickly, and expose much of the combustible tissue to the air. Consequently, grass fires spread rapidly but burn out quickly (Anderson 1982). Wind strongly affects fire and the most extreme rates of spread can be as high as 30% of the wind speed.

Grass is found in most ecosystems, but is the dominant and most influential fuel in arid and range lands. Southern ecosystems with dominant grass fuels are open longleaf pine forests, dry prairie, the Everglades, and pastures. When surveying an area to determine fuel model, the percentage of grass should exceed 66%.

Grass fuel models make up FPBS models 1-3.

Shrub layers that are three to six feet high present a deep fuel bed, with fine fuels (leaves, litter) exposed throughout the fuel bed, leading to potential for long flame lengths and high energy release during fires. Most often wind does not affect fire spread of shrub fuels - unless stronger than 8 mph at mid-flame height. Shrubs are found in most ecosystems. In pine-flatwoods and pocosins on the Coastal Plain, palmetto, wax myrtle, and gallberry are the dominant shrubs. They contain oils and other volatiles that burn readily, and they often form a dense understory and midstory, making up 35-40% of total cover. In the upper Piedmont and mountains, rhododendron and mountain laurel are common shrubs that also increase fire intensity.

Shrub fuel models make up FPBS models 4-7.

Timber fuel models are actually models of "litter" on the forest floor. Timber litter consists of the fallen leaves, needles, small woody debris and larger logs on the forest floor. Timber litter occurs in all ecosystems dominated by trees, but it is the main surface fuel in closed canopy forest systems, such as short-needle conifers or hardwoods where understory growth is limited (Anderson 1982). Compact layering of the fuels on the forest floor reduces the surface area exposed to air and therefore the rate of combustion and fire spread within the fuels tends to be slow.

Timber fuel models make up FPBS models 8-10.

Logging slash includes the residual material after timber harvesting or silviculture operations like thinning, pruning, or removing competitive understory plants. This debris includes branches, bark, logs, and understory shrubs or trees. Large diameter logs and branches have low surface area to volume ratios, do not readily ignite, and fires move through them relatively slowly unless there is also an accumulation of fine fuels such as needles and small branches. Although large (heavy) fuels ignite slowly, once they are burning they are difficult to extinquish and can release large amounts of heat.

Slash fuel models make up FPBS models 11-13.

The following thirteen fuel models, known variously as "standard", "stylized", "Fire Behavior" and "NFFL" models, were were originally tabulated by Rothermel (1972) and later expanded by Albini (1976) . These fuel models are a complete set of fuel inputs needed to use the Rothermel fire spread model (load and surface-are-to-volume ratio by size class and component, extinction moisture, heat content by component, and fuelbed depth) used in BEHAVE Plus and other fire behavior prediction systems to predict real time fire behavior across a range of conditions.

These models are conceptually organized into grass, shrub, timber, and slash fuel types. Below, the fuel model number from 1-13, example systems from the southeast, and related fuel models from the NFDRS model system are presented. The fuel loading and fuel bed depth for these 13 fuel types are shown in Table: FBPS Fuel Models.

Grass fuel models

1. grasslands and savannas (correlates to NFDRS models A, L, and S)
2. open shrub land, pine and scrub oak stands covering less than 2/3 area (correlates to NFDRS models C and T)
3. tall prairie and marshland grasses where influence of wind is high (correlates to NFDRS model N)

Shrub fuel models

4. stands of mature shrubs, closed jack pine stands (correlates to NFDRS models B and O)
5. young green stands with no dead wood such as laurel or vine maple (no model represented)
6. intermediate shrub stands and cured hardwood slash (correlates to NFDRS models F and Q)
7. stands of shrubs 2 - 6 feet such as palmetto-gallberry with pine overstory (correlates to NFDRS model D)

Timber fuel models

8. closed canopy stands of short-needle conifers or hardwoods that have leafed out and support fire in the compact litter layer (correlates to NFDRS models H and R)
9. long-needle conifer and hardwood stands (correlates to NFDRS models E, P, and U)
10. any stand with large quantities of dead-down fuel (correlates to NFDRS model G)

Slash fuel models

11. conifer or hardwood stands with light partial cuts or thinning (correlates to NFDRS model K)
12. heavily thinned conifer stands, clearcuts, medium – heavy partial cuts (correlates to NFDRS model J)
13. clearcuts and heavy partial cuts in mature stands where slash is dominated by material with diameter exceeding 3 inches (correlates to NFDRS model I)

The twenty National Fire Danger Rating System (NFDRS) fuel models, similar to the FBPS fuel models but with extra parameters needed for fire danger prediction, are typically used in fire danger rating systems such as the National Fire Danger Rating System (NFDRS). These 20 fuel types were selected for modeling based on: their significance as a fuel type, interest expressed by fire managers, and the availability of fuels and fire behavior data with which to develop and test the models. Since the target area for a fire danger rating system often encompasses tens of thousands of acres the fuel models need only reflect the “typical” fuel situation in each fire-danger rating area (Deeming 1975). A fuel model key also exists which helps to determine which fuel models are appropriate for a particular area (Deeming 1977). The USDA Forest Service also provides a map of the NFDRS fuel models across the lower 48 states at 1 km resolution.

NFDRS Fuel Models

1. western annual grasses
2. California mixed chaparral
3. pine grass savanna
4. southern rough
5. hardwoods (winter)
6. intermediate brush
7. short needle pine (heavy dead)
8. short needle pine (normal dead)
9. heavy logging slash
10. intermediate logging slash
11. light logging slash
12. western perennial grass
13. sawgrass
14. high pocosin
15. southern pine plantation
16. Alaskan black spruce
17. hardwoods (summer)
18. tundra
19. sagebrush-grass
20. western long-needled conifer

The Fuelbed Characteristic Classification System (or FCCS; Sandberg et al. 2001) is comprised of a large database of physical fuel bed parameters, an expert system to select fuel beds, and a system for estimating fire potential. Currently under development, FCCS is designed to provide the best possible fuel estimates and potential fire parameters based on as much or as little site-specific information as is available. For example, FCCS will provide detailed fuel bed estimates based on either specific fuel data or general site data that are available for Ecoregion Divisions, vegetation forms, cover types, or other data obtained from remote sensing, forest inventories, models, etc... The detailed fuelbed estimates can then be used to support fire hazard assessments and make fuel treatment decisions.

The prototype FCCS database includes 200+ fuelbeds common to North America, and a process has been established to add several tens of thousands more fuelbeds in the next several years. The database describes the size, abundance, physical character, and arrangement of a wildland fuelbed. Fuel beds are compartmentalized by vertical position:

• forest canopy,

• shrub vegetation,

• low vegetation,

• woody fuels,

• litter, and

• moss/duff strata.

An expert system allows a user to select a fuelbed from general site data and to adjust fuelbeds in the database based on specific site data or other available information. This component of the system also calculates and summarizes fuelbed properties by vertical strata, and uses a set of look-up tables to assign properties based on vegetation species or physical characteristics. Fuel bed fire potentials (on scales of 0-10) are estimated for surface fire behavior, crowning potential, and fuel consumption. These potentials are calculated from the loading, heat content, bulk density, and characteristic thickness of fuel elements without consideration of moisture content or environmental conditions. They represent the potential fire behavior and effects of an oven-dry fuelbed with no wind or slope influence.

In addition to reporting fire potentials on numeric scales from low to high, fuel bed classifications are currently being designed to be compatible with several existing fire and landscape assessment models. It is anticipated that future work will allow these fuel bed models to interface with other fire models such as Behave and FARSITE.

Fire danger describes how easy it is to ignite vegetation, how difficult a fire may be to control, and how much damage a fire may do. Fire danger rating systems produce qualitative and/or numeric indices of fire potential based on fuels, topography and weather. These rating systems allow fire managers to estimate present and future fire danger for a given area. They are commonly used as guides for initiating pre-suppression activities and selecting the appropriate level of initial response to a reported wildfire where detailed, site and time-specific information is unavailable. Usually these systems require mathematical descriptions of fuel models and their respective fuel properties as input.

The National Fire Danger Rating System (NFDRS), started in the 1960s, is the most commonly used system in the U.S. for rating fire danger. The NFDRS is a set of computer programs and algorithms that allow land management agencies to estimate present and future fire danger for a given area based on fuels, topography and weather. NFDRS characterizes fire danger by evaluating the upper limit of fire behavior in an area during a 24-hour period. Each day during the fire season, spot measurements of fire danger, calculated using the NFDRS at specific weather stations, are interpolated and mapped on a national basis by the Wildland Fire Assessment System (WFAS-MAPS) in Missoula, Montana. Two maps are produced daily during the fire season, an observed fire danger map and a forecasted fire danger map. The NFDRS outputs give relative ratings of the potential growth and behavior of any wildfire. Fire danger ratings are guides for initiating pre-suppression activities and selecting the appropriate level of initial response to a reported wildfire where detailed, site and time-specific information is unavailable. See: History of NFDRS.

The Oklahoma Fire Danger Model, operational since 1996, is a prototype next-generation model of the NFDRS. In contrast to the NFDRS, which uses one weather reading per day, the OFDM uses a real-time automated weather station network (the Oklahoma Mesonet) to provide fire danger ratings at a smaller resolution than the NFDRS.

The Canadian Forest Fire Danger Rating System (CFFDRS), under development since 1968, is a national system for rating the risk of forest fires in Canada. This system is made up of two subsystems–the Canadian Forest Fire Weather Index (FWI) System and the Canadian Forest Fire Behavior Prediction (FBP) System. Similar to the NFDRS, the Canadian system publishes daily maps of fire danger.

In 1954 there were eight different fire-danger rating systems in use across the country. Fire control conferences called by the United States Forest Service highlighted need for a nationwide system. Improved communication and transportation systems made mutual assistance agreements between fire control agencies more practical than in the past. State compacts, and in the case of the Federal government, interagency and interregional agreements were bringing fire control teams together from widely separated areas of the country. The increased cooperation required a national system for estimating fire danger and fire behavior to improve and simplify communications among all people concerned with wildland fires.

Work on a national rating system began in 1959. By 1961, the basic structure for a four-phase rating system (ignition, risk, fuel energy, and fire spread) had been outlined and the fire spread phase was ready for field testing. Adaptations, interpretations, and additions to the fire spread phase quickly followed, making it obvious that the fire spread phase was not uniformly applicable across the country.

A new research project was initiated in 1965 with the goal of developing a complete, comprehensive, National Fire-Danger Rating System by 1972. In 1970, a preliminary version of the system was field tested in Arizona and New Mexico. In 1971, an improved version of the system was used operationally in the Southwest. Field trials were also conducted elsewhere across the country at stations from Maine to California and from Florida to Alaska. The National Fire Danger Rating System (NFDRS) was implemented in 1973, and then revised during a three year project from 1975-1978 and reissued as the 1978 NFDRS (NOAA 2003).

The 1978 NFDRS

1. The system only considered “initiating fire” which is fire that is not behaving erratically and does not have spotting or crowning.

2. Ratings of difficulty of containment were based solely on fire behavior. Other factors affecting difficulty of containment (accessibility, soil and resistance to line construction, and so on) were to be assessed by other means.

3. Fire behavior, defined in part by flame length at the fire front, was assumed to be directly related to the difficulty of containment.

4. Fire danger rating were based on the “worst” fire conditions by using meteorological measurements when fire behavior is normally the most extreme (i.e., in early afternoon, sites in the open, and sites of southerly or westerly exposure).

5. The system provided ratings that were physically interpretable in terms of fire occurrence and behavior. These ratings could then be used alone or in combination.

6. The ratings were relative, not absolute.

• modified responses to drought, with open ended rather than closed index scales, to reflect effect of changing day length on burning conditions and fuel moisture
• separate human and lightning caused occurrence indices
• expanded ignition component and a spread component
• increased number of fuel models from 9 to 20, increased number of slope definitions from 3 to 5, and predictive models for the moisture content of live grasses, forbs, and woody shrubs.

The updated 1978 NFDRS system contained 3 fire behavior components, 5 fuel classes (3 dead, 2 live), and the following 4 indices for rating fire danger:

1. human caused fire occurrence index: this index is derived from an assessment of human caused fire sources in the rating area, and the ignition component which is the likelihood that a firebrand will cause a reportable fire.

2. lightning caused fire occurrence index: this index was derived from the ignition component and lightning risk, an indicator of thunderstorm and lightning activity.

3. burning index (BI): this index was derived from the spread component (a relative index of rate of fire spread), the energy release component (a relative index of the amount of heat released per unit area in the flaming zone of an initiating fire), the live and dead fuel moisture, and the fuel model. The BI indicates the difficulty of containment of a single fire.

4. fire load index (FLI): the BI and the expected number of fires predicted by the human caused and lightning caused fire occurrence indices combined to produce the fire load index which was a measure of the total fire containment job. The FLI was the NFDRS cumulative index integrating risk, ignition probability, and fire behavior potential.

The 1978 system was updated again in 1988 to correct for some deficiencies when applying the model in the eastern United States. These updates improved the capability of the model to respond to drought in humid environments, provided flexibility to reflect the greening and curing of live fuels, corrected the problems of overrating fire danger in the autumn or after rainfall, and adjusted the fuel models to better predict fire danger in humid climates (Burgan 1988).

Although all the original 1978 indices provide specific types of information, the main indices currently used for fire danger rating are the spread component (SC), energy release component (ERC) and burning index (BI). The KBDI has also been incorporated into the NFDRS.

By the summer of 1976 data from more than 800 fire-weather stations across the US were being processed by a central mainframe computer (AFFIRMS), and the system indices and components were calculated each day for over 400 stations. The original computer program for calculating NFDRS has been updated several times. The NFDRSPC, or national fire danger rating system for personal computers, was developed in 1988 to make adjustment for those managers in eastern states who were using the 1988 version of the 1978 NFDRS. This system was essentially a revision to the AFFIRMS computer system with a window/menu-oriented user interface (Donaldson et al. 1990). Since then, PCDANGER, a menu-driven PC- based application has replaced these programs. This new program is beneficial for managers because it provides local flexibility in computing and graphing fire-danger-rating indices and can compute fire danger indices for both the 1978 and 1988 versions of the NFDRS (Bradshaw et al. 1997).