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.
3-5 Figure 3-3. Temperature-Dewpoint Spread Effect on Relative Humidity
Relative humidity (RH) depends on the temperature-dewpoint spread. In Figure 3-3, dewpoint is constant but temperature decreases from left to right. On the left panel, relative humidity is 50 percent, which indicates the air parcel could hold twice as much water vapor as is actually present. As the air parcel cools, the temperature-dewpoint spread decreases while relative humidity increases. When the air parcel’s temperature cools to equal its dewpoint (11 °C), its capacity to hold water vapor is reduced to the amount actually present.
The temperature-dewpoint spread is zero, relative humidity is 100 percent, and the air parcel is now saturated.
3.7 Change of Phase. Water changes from one state of matter (solid, liquid, or vapor) to another at the temperatures and pressures experienced near the surface of the Earth.
Interestingly, water is the only substance on Earth that exists naturally in all three phases:
as water droplets and ice crystals (visible as clouds) and as water vapor.
Water has some unique thermal properties which make it a powerful heat transport mechanism. It has the highest specific heat capacity of any naturally occurring substance (see Table 2-4). That means water has a much higher capacity for storing heat energy (with little resulting temperature change) than other substances. These properties make water an ideal heat transport mechanism, and have important implications on weather and climate.
3.7.1 Latent Heat. Latent heat is the quantity of heat energy either released or absorbed by a unit mass of a substance when it undergoes a phase transition (change of state).
Units are typically expressed in terms of joules per gram (J/g). Figure 3-4 below illustrates the latent heat transactions that occur when water undergoes phase transition.
3-6 Figure 3-4. Latent Heat Transactions when Water Undergoes Phase Transition
Heat is exchanged between water and its environment during phase transition. Although the temperature of the environment changes in response, the temperature of the water undergoing the phase transition remains constant until the phase change is complete; that is, the available heat, latent heat, is involved exclusively in changing the phase of water and not in changing its temperature. There are six phase transitions, three of which are associated with the absorption of latent heat by water from the environment (melting, evaporation, and sublimation), and three of which are associated with the release of heat energy by water to the environment (freezing, condensation, and deposition).
Melting is the phase transition by which a solid is changed to a liquid. During melting, water absorbs 334 joules per gram due to the latent heat of fusion. Freezing, the reverse process, releases 334 joules per gram back to the environment.
Evaporation is the phase transition by which a liquid is changed to a vapor. During evaporation, water absorbs 2,501 joules per gram due to the latent heat of vaporization.
Condensation, the reverse process, releases 2,501 Joules per gram back to the environment.
Sublimation is the phase transition by which a solid is changed to a vapor. During sublimation, water absorbs 2,834 joules per gram due to the latent heat of sublimation.
Deposition, the reverse process, releases 2,834 joules per gram back to the environment.
3-7 Table 3-1. Latent Heat of Water at 0 °C Latent Heat Type Energy Exchange (J/g) Latent heat of sublimation 2,834 Latent heat of vaporization 2,501 Latent heat of fusion 334 The amount of energy associated with latent heat exchange should not be understated.
An average hurricane releases 52 million trillion (5.2 x 1019) joules per day as water vapor condenses into clouds and precipitation. This is equivalent to about 40 times the total worldwide energy consumption per day in 2005!
4-1 CHAPTER 4. EARTH-ATMOSPHERE HEAT IMBALANCES
4.1 Introduction. Weather is not a capricious act of nature, but rather the atmosphere’s response to unequal rates of radiational heating and cooling across the surface of the Earth and within its atmosphere. The absorption of incoming solar radiation causes heating, while the emission of outgoing terrestrial radiation causes cooling. However, imbalances in the rate of heating and cooling create temperature gradients.2 Atmospheric circulations and weather are the atmosphere’s never-ending attempt to redistribute this heat and achieve equilibrium.
4.2 The Earth-Atmosphere Energy Balance. The Earth-atmosphere energy balance is the balance between incoming energy from the sun (solar radiation) and outgoing energy from the Earth (terrestrial radiation), as seen in Figure 4-1. When solar radiation reaches the Earth, some is reflected back to space by air (8 percent), clouds (17 percent), or the surface (6 percent). Some is absorbed by water vapor/dust/ozone (19 percent) or by clouds (4 percent). The remainder is absorbed by the Earth’s surface (46 percent).
Figure 4-1. Earth-Atmosphere Energy Balance
100 units of incoming radiation from the sun is balanced by 100 units of outgoing radiation from the Earth.
2 A gradient describes the rate of change of a feature (in this case temperature) per unit of distance.