2-6 Figure 2-4. Solar Zenith Angle
2.6.2 Conduction. Conduction is the transfer of energy (including heat) by molecular activity from one substance to another in contact, or through, a substance. Heat always flows from the warmer substance to the colder substance. The rate of heat transfer is greater with larger temperature differences and depends directly on the ability of the substance(s) to conduct heat. During conduction, the warmer substance cools and loses heat energy, while the cooler substance warms and gains heat energy.
Heat (thermal) conductivity is the property of a substance that indicates its ability to conduct heat as a consequence of molecular motion. Units are Watts per meter-Kelvin (W m-1 K-1). Table 2-3 below provides the heat (thermal) conductivity of various substances. Note that air is a poor thermal conductor.
2-7 Table 2-3. Heat (Thermal) Conductivity of Various Substances Material Phase Heat (Thermal) Conductivity (W m-1 K-1) Silver Solid 429 Copper Solid 401 Aluminum Solid 250 Iron Solid 80 Sand (saturated) Solid 2.7 Water (ice) Solid (0 °C) 2.18 Sandstone Solid 1.7 Limestone Solid 1.26 – 1.33 Glass Solid 1.05 Water (liquid) Liquid 0.58 Sand (dry) Solid 0.35 Soil Solid 0.17 – 1.13 Wood (oak) Solid 0.17 Wood (balsa) Solid 0.055 Snow Solid (<0 °C) 0.05 – 0.25 Air Gas 0.024 Water (steam) Gas (125 °C) 0.016 All measurements are at 25 °C unless otherwise noted.
Note: 1 K equals -272.15 °C.
2.6.3 Convection. Convection is the transport of heat within a fluid, such as air or water, via motions of the fluid itself. This type of heat flow takes place in liquids and gases because they can move freely, and it is possible to set up currents within them. Water boiling in a pot is an example of convection. Because air is a poor thermal conductor, convection
2-8 plays a vital role in the Earth’s atmospheric heat transfer process. Figure 2-5 illustrates examples of various heat transfer processes.
Figure 2-5. Heat Transfer Examples1
2.7 Thermal Response. Whether by radiation, conduction, convection, or a combination of these, the temperature response to the input (or output) of some specified quantity of heat varies from one substance to another. Specific heat capacity, also known simply as specific heat, is defined as the measure of heat energy required to increase the temperature of a unit quantity of a substance by a certain temperature interval. Specific heat capacity is typically expressed in units of joules per gram-Kelvin (J g-1 K-1). Thus, two different substances with identical temperature measurements do not necessarily possess the same amount of heat energy. When exposed to the same amount of heat energy, a substance with a low specific heat capacity warms up more than a substance with a higher specific heat capacity.
1 The source of this and other material labeled COMET® is the Web site at http://meted.ucar.edu/ of the University Corporation for Atmospheric Research (UCAR), sponsored in part through cooperative agreement(s) with the National Oceanic and Atmospheric Administration (NOAA), U.S. Department of Commerce (DOC). ©1997-2015 University Corporation for Atmospheric Research. All Rights Reserved.
2-9 Table 2-4. Specific Heat Capacity of Various Substances Substance Phase Specific Heat Capacity (J g-1 K-1) Water (steam) Gas (100 °C) 4.22 Water Liquid (25 °C) 4.18 Wood (balsa) Solid 2.90 Water (ice) Solid (0 °C) 2.05 Wood (oak) Solid 2.00 Soil (wet) Solid 1.48 Sandy clay Solid 1.38 Air (sea level, dry) Gas 1.01 Asphalt Solid 0.92 Clay Solid 0.92 Aluminum Solid 0.91 Brick (common) Solid 0.90 Concrete Solid 0.88 Glass Solid 0.84 Limestone Solid 0.84 Sand (quartz) Solid 0.83 Soil (dry) Solid 0.80 Granite Solid 0.79 Iron Solid 0.46 Copper Solid 0.39 Mercury Liquid 0.14 Lead Solid 0.13 All measurements are at 25 °C unless otherwise noted.
Note: 1 K equals -272.15 °C.
2-10 \Water has the highest specific heat capacity of any naturally occurring substance. That means it has a much higher capacity for storing heat energy than other substances, such as soil, sand, rock, or air. Water can store large amounts of heat energy while only experiencing a small temperature change.
Figure 2-6 below compares the specific heat capacity of water and sand. The specific heat capacity of water is more than five times that of quartz sand. Thus, 4.18 joules of heat are required to raise the temperature of one gram of water by 1 °C, while only 0.83 joules are required to raise the temperature of one gram of quartz sand by 1 °C. This is one reason why beach sand is hotter than water on a sunny, summer afternoon.
Figure 2-6. Specific Heat Capacity: Water Versus Sand
The difference in specific heat capacities is one of the primary reasons why the temperature of a body of water, such as a lake or the ocean, is less variable with time than the surface temperature of land. Water heats up more slowly than land during the day and during summer, and cools down more slowly at night and during the winter. Thus, a body of water exhibits greater resistance to temperature change, called thermal inertia, than does a land mass.
Heat flow differences are another reason why water bodies warm up and cool down more slowly than land. Incoming solar radiation penetrates water to significant depths, but can only heat the top skin layer of soil and rock. Also, since water is a fluid, its heat energy can be circulated through great volumes and depths via convection. Water temperature changes occur to depths of six meters (20 feet) or more on a daily basis, and 200 to 600 meters (650 to 1950 feet) annually. The process is more problematic over land since heat must be transferred via the slow process of conduction. Land temperature changes occur to depths of only 10 centimeters (4 inches) on a daily basis and 15 meters (50 feet) or less annually.
2-11 Water is much more resistant to temperature changes than land. It warms up and cools down more slowly than land and helps to moderate nearby air temperature. This is why islands and localities located immediately downwind from the ocean or a large lake (maritime locations) exhibit smaller diurnal and seasonal temperature variations than localities well inland (continental locations). Figure 2-7 illustrates this effect. Although both cities are at approximately the same latitude, the temperature is much less variable at San Francisco (maritime) than Kansas City (continental).
Figure 2-7. Variation of Mean Daily Temperatures for San Francisco (Maritime) and Kansas City (Continental)
2.8 Temperature Variations with Altitude. A lapse rate of temperature is defined as a decrease in temperature with height. In Table 1-2, it was stated that the temperature decreases 6.5 °C/kilometer (3.57 °F/1,000 feet) in the standard atmosphere. But since this is an average, the exact value seldom exists. In fact, temperature in the troposphere sometimes remains constant or even increases with height. Caution should be taken when using the standard lapse rate to estimate the freezing level. Quite often the boundary layer is dry adiabatic and the estimate of freezing level could be in error.
2-12
2.8.1 Atmospheric Sounding. An atmospheric sounding, or simply sounding, is a plot of the vertical profile of one or more atmospheric parameters, such as temperature, dewpoint, or wind above a fixed location. Soundings are used extensively by meteorologists to determine the state of the atmosphere.
2.8.2 Isothermal Layer. An isothermal layer is a layer within the atmosphere where the temperature remains constant with height (see Figure 2-8).
Figure 2-8. Sounding with an Isothermal Layer
2.8.3 Temperature Inversion. A temperature inversion, or simply inversion, is a layer in which the temperature increases with altitude. If the base of the inversion is at the surface, it is termed a surface-based inversion. If the base of the inversion is not at the surface, it is termed an inversion aloft (see Figure 2-9).
A surface-based inversion typically develops over land on clear nights when wind is light. The ground radiates and cools much faster than the overlying air. Air in contact with the ground becomes cool, while the temperature a few hundred feet above changes very little. Thus, temperature increases with height.
An inversion may also occur at any altitude when conditions are favorable. For example, a current of warm air aloft overrunning cold air near the surface produces an inversion aloft. Inversions are common in the stratosphere.
The principal characteristic of an inversion layer is its marked stability, so that very little turbulence can occur within it. Turbulence will be discussed at length in Chapter 17.
Figure 2-9. Sounding with a Temperature Inversion
KM FEET 200 - 39,000 11735,000 10- 300 9- - 30,000 ALTITUDE (U.S. STANDARD ATMOSPHERE) 8- PRESSURE (MILLIBARS) - 25,000 400 7 6--20,000 500 5-
- 15,000 600 4
INVERSION (ALOFT) 700 3- - 10,000 800 2 5,000 900
INVERSION 1;050 (SURFACE BASED) 0- - o -60 -50 -40 -30 -20 -10 10 20 30 40°C -20 20 40 60
TEMPERATURE 2-13
3-1 CHAPTER 3. WATER VAPOR
3.1 Introduction. Water vapor is the gaseous form of water and one of the most important of all constituents of the atmosphere. It constitutes only a small percentage of the Earth’s atmosphere, varying from only trace amounts to 4 percent by volume, and its amount varies widely in space and time. Approximately half of all of the atmospheric water vapor is found below 2 kilometers (6,500 feet) altitude, and only a minute fraction of the total occurs above the tropopause.
Water vapor is important, not only as the raw material for clouds and precipitation
(e.g., rain and snow), but also as a vehicle for the transfer of heat energy and as a regulator of the Earth’s temperatures through absorption and emission of radiation, most significantly in the thermal infrared (i.e., the greenhouse effect). The amount of water vapor present in a given air sample may be measured in a number of different ways, involving such concepts as relative humidity and dewpoint. Before we talk about these subjects, we will first discuss how water cycles through the Earth-atmosphere system.
3.2 The Hydrologic Cycle. The hydrologic cycle (see Figure 3-1) involves the continuous circulation of water in the Earth-atmosphere system. Water vapor plays a critical role in the cycle.
Figure 3-1. The Hydrologic Cycle
3-2
3.2.1 Evaporation. Evaporation is the phase transition by which a liquid is changed to a vapor (gas). In meteorology, the substance we are concerned about the most is water, and the primary source is the ocean. On average, about 120 centimeters (47 inches) is evaporated into the atmosphere from the ocean each year. For evaporation to take place, energy is required. The energy can come from any source: the sun, the atmosphere, the Earth, or objects on the Earth, such as humans.
Everyone has experienced evaporation personally. When the body heats up due to the air temperature, or through exercise, the body sweats, secreting water onto the skin.
The purpose is to cause the body to use its heat to evaporate the liquid, thereby removing heat and cooling the body. The same effect can be seen when you step out of a shower or swimming pool. The coolness you feel is from the removal of bodily heat used to evaporate the water on your skin.
3.2.2 Transpiration. Transpiration is the evaporation of water from plants. In most plants, transpiration is a passive process largely controlled by the humidity of the atmosphere and the moisture content of the soil. Of the transpired water passing through a plant, only 1 percent is used in the growth process of the plant. The remaining 99 percent is passed into the atmosphere.
3.2.3 Sublimation. Sublimation is the phase transition by which a solid is changed into vapor (a gas) without passing through the liquid phase. In the atmosphere, sublimation of water occurs when ice and snow (solids) change into water vapor (a gas).
3.2.4 Condensation. Condensation is the phase transition by which vapor (a gas) is changed into a liquid. In the atmosphere, condensation may appear as clouds, fog, mist, dew, or frost, depending upon the physical conditions of the atmosphere.
3.2.5 Transportation. Transportation is the movement of solid, liquid, and gaseous water through the atmosphere. Without this movement, the water evaporated over the ocean would not precipitate over land.
3.2.6 Precipitation. Precipitation results when tiny condensation particles grow through collision and coalescence.
3.2.7 Runoff. Runoff occurs when there is excessive precipitation and the ground is saturated
(i.e., cannot absorb any more water). This runoff flows into streams and rivers and eventually back into the sea.
Evaporation of this runoff into the atmosphere begins the hydrologic cycle over again.
Some of the water percolates into the soil and into the ground water, only to be drawn into plants again for transpiration to take place.
3.2.8 Infiltration. Infiltration is the movement of water into the ground from the surface.
3-3
3.2.9 Groundwater Flow. Groundwater flow is the flow of water underground in aquifers.
The water may return to the surface in springs or eventually seep into the oceans.
3.2.10 Plant Uptake. Plant uptake is water taken from the groundwater flow and soil moisture.
3.3 Saturation. Saturation is the maximum possible quantity of water vapor that an air parcel can hold at any given temperature and pressure. The term saturated air means an air parcel has all the water vapor it can hold, while unsaturated air means an air parcel can hold more water vapor.
3.4 Relative Humidity. Relative humidity is the ratio, usually expressed as a percentage, of water vapor actually in the air parcel compared to the amount of water vapor the air parcel could hold at a particular temperature and pressure.
𝑅𝑅𝑅𝑅𝑅𝑅𝑅𝑅 𝐻𝐻𝐻𝑅𝐻𝑅𝑅𝑦 = 𝑊𝑅𝑅𝑅𝑊 𝑅𝑅𝑣𝑣𝑊 𝑐𝑣𝑐𝑅𝑅𝑐𝑅𝑊𝑅𝑅𝑅𝑊 𝑅𝑅𝑣𝑣𝑊 𝑐𝑅𝑣𝑅𝑐𝑅𝑅𝑦 While relative humidity is the most common method of describing atmospheric moisture, it is also the most misunderstood. Relative humidity can be confusing because it does not indicate the actual water vapor content of the air, but rather how close the air is to saturation. An air parcel with 100 percent relative humidity is saturated, while an air parcel with relative humidity less than 100 percent is unsaturated.
An air parcel’s capacity to hold water vapor (at a constant pressure) is directly related to its temperature. It is possible to change an air parcel’s relative humidity without changing its water vapor content. Figure 3-2 illustrates the concept below. An air parcel at sea level at a temperature of 30 °C has the capacity to hold 27 grams of water vapor.
If it actually held 8 grams, its relative humidity would be 30 percent, and it would be unsaturated. However, if the air parcel’s temperature decreases to 20 °C, its water vapor storage capacity decreases to 15 grams and its relative humidity rises to 53 percent.
At 10 °C, the air parcel’s water vapor storage capacity decreases to equal the amount of water vapor it actually holds (8 grams), its relative humidity increases to 100 percent, and it becomes saturated. During this cooling process, the air parcel’s actual water vapor content remained constant, but relative humidity increased with decreasing temperature.
3-4 Figure 3-2. Temperature Effects on Relative Humidity
3.5 Dewpoint. Dewpoint is the temperature an air parcel must be cooled at constant pressure and constant water vapor pressure to allow the water vapor in the parcel to condense into water (dew). When this temperature is below 0 °C (32 °F), it is sometimes called the frost point. Lowering an air parcel’s temperature reduces its capacity to hold water vapor.
3.6 Temperature-Dewpoint Spread (Dewpoint Depression). The difference between an air parcel’s temperature and its dewpoint is the dewpoint depression, or commonly referred to as the spread. Surface aviation weather reports (e.g., Aviation Routine Weather Reports (METAR)/Aviation Selected Special Weather Reports (SPECI)) provide observations of both temperature and dewpoint. The temperature greatly affects the air parcel’s ability to hold water vapor, while the dewpoint indicates the actual quantity of water vapor in the parcel. As the spread decreases, relative humidity increases. When the spread decreases to zero, relative humidity is 100 percent, and the air parcel is saturated. Figure 3-3 below illustrates the relationship between temperaturedewpoint spread and relative humidity.
Surface temperature-dewpoint spread is important in anticipating fog, but has little bearing on precipitation. To support precipitation, air must be saturated through thick layers aloft.