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.
4-2 However, since the Earth is much cooler than the sun, its radiating energy is much weaker (long wavelength) infrared energy. We can indirectly see this energy radiate into the atmosphere as heat (e.g., rising from a hot road, creating shimmers on hot sunny days). The Earth-atmosphere energy balance is achieved as the energy received from the sun (solar radiation) balances the energy lost by the Earth back into space (terrestrial radiation). In this way, the Earth maintains a stable average temperature.
The absorption of infrared radiation trying to escape from the Earth back to space is particularly important to the global energy balance. Energy absorption by the atmosphere stores more energy near its surface than it would if there was no atmosphere.
The average surface temperature of the moon, which has no atmosphere, is -18 °C (0 °F).
By contrast, the average surface temperature of the Earth is 15 °C (59 °F). This heating effect is called the greenhouse effect.
Greenhouse warming is enhanced during nights when the sky is overcast (see Figure 4-2).
Heat energy from the Earth can be trapped by clouds, leading to higher temperatures as compared to nights with clear skies. The air is not allowed to cool as much with overcast skies. Under partly cloudy skies, some heat is allowed to escape, and some remains trapped. Clear skies allow for the most cooling to take place.
Figure 4-2. Greenhouse Effect on Nighttime Radiational Cooling
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4.3 Heat Imbalances Between Earth’s Surface and the Atmosphere.
The Earth-atmosphere energy balance numbers (see Figure 4-1) indicate that both sensible heat (7 percent) and latent heat (24 percent) processes transfer heat from the Earth’s surface into its atmosphere. Both processes are necessary to prevent the Earth’s surface from continually heating up and the atmosphere from continually cooling down.
4.3.1 Sensible Heating. Sensible heating involves both conduction and convection. It occurs due to differences in air density. Warm air is less dense than cool air.
On warm sunny days, the Earth’s surface is heated by incoming solar radiation or insolation. However, the heating is somewhat uneven because certain areas of the Earth’s surface absorb more heat from the sun than others. Heat is conducted from the relatively warm ground to the cooler overlying air, which warms a shallow layer of air near the ground. The heated air expands, becomes less dense than the surrounding cooler air, and rises. Through this process, a large bubble of warm air called a thermal rises and transfers heat energy upwards (see Figure 4-3). Cooler, denser air sinks toward the ground to replace the rising air. This cooler air becomes heated in turn, rises, and repeats the cycle.
Figure 4-3. Development of a Thermal
In this manner, convection transports heat from the Earth’s surface into the atmosphere.
Because air is a poor conductor of heat (see Table 2-3), convection is much more important than conduction as a heat transport mechanism within the atmosphere.
4-4 Figure 4-4. Example of Convection in the Atmosphere
4.3.2 Latent Heat. The phase transition of water and associated latent heat exchanges (see subparagraph 3.7.1) are largely responsible for transferring the excess heat from the surface of the Earth into its atmosphere. As the Earth’s surface absorbs radiation, some of the heat produced is used to evaporate (vaporize) water from oceans, lakes, rivers, soil, and vegetation. The water absorbs heat energy due to the latent heat of vaporization.
Some of this water vapor condenses to microscopic water droplets or deposits as ice crystals that are visible as clouds. During cloud formation, the water vapor changes state, and latent heat is released into the atmosphere. During this process, the excess heat is transferred from the Earth’s surface into its atmosphere.
4.4 Heat Imbalance Variations with Latitude. Global imbalances in radiational heating and cooling occur not only vertically between the Earth’s surface and its atmosphere, but also horizontally with latitude. Since the Earth is essentially spherical, parallel beams of incoming solar radiation strike lower latitudes more directly than higher latitudes (see Figure 4-5); that is, the solar zenith angle is lower, and the sun is more directly overhead in equatorial regions than at the poles. At higher latitudes, solar radiation is spread over a larger area and is less intense per unit surface area than at lower latitudes.
Thus, the Earth absorbs more solar radiation at lower latitudes than higher latitudes, which creates heat imbalances and temperature gradients between the Equator and the poles.
4-5 Figure 4-5. Solar Zenith Angle Variations with Latitude
The emission of terrestrial radiation also varies by latitude, but less so than the absorption of solar radiation. Terrestrial radiation emission decreases with increasing latitude due to a drop in temperature with latitude. Thus, at higher latitudes, the annual rate of cooling exceeds the rate of warming, while the reverse is true at lower latitudes.
Averaged over all latitudes, incoming solar radiation must equal outgoing terrestrial radiation. Otherwise, the Earth would be unable to maintain a constant average temperature. About 35° latitude in both hemispheres is where incoming and outgoing radiation is equal. This implies there is annual net cooling at higher latitudes and net warming at lower latitudes. However, we know this is untrue. The excess heat in the tropics must be transported polar by some mechanism(s). This poleward heat transport is accomplished by atmospheric circulations, weather, and ocean currents.
4.5 Seasons. Seasons are caused by the tilt of the Earth’s rotational axis as the Earth orbits the sun (see Figure 4-6). The Earth’s rotational axis is tilted by 23½° from the perpendicular drawn to the plane of the Earth’s orbit about the sun and points the same direction in space all year long. The North Pole is tilted most directly toward the sun Solstice (~December 22). Thus, in the Northern Hemisphere, the longest day of the year
4-6 (lowest solar zenith angle) occurs on the Summer Solstice, while the shortest day of the year (highest solar zenith angle) occurs on the Winter Solstice. Day and night are of equal length (12 hours) worldwide on the Vernal Equinox (~March 21) and the Autumnal Equinox (~September 23).
Figure 4-6. Solar Zenith Angle Variations with Northern Hemisphere Seasons
Figure 4-7 illustrates the average seasonal temperature variation in the Northern Hemisphere for both maritime and continental locations. Note that the warmest (coldest days) of the year occur after the summer (winter) solstice. This is due to the time lag necessary for heat flow processes to fully heat (cool) the surface of the Earth.
Figure 4-7. Average Seasonal Temperature Variation in the Northern Hemisphere
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4.6 Diurnal Temperature Variation. Diurnal temperature variation is the daytime maximum and nighttime minimum of air temperature due to variations of insolation caused by the rising and setting of the sun (variations of solar zenith angle) as the Earth rotates around its axis. Figure 4-8 depicts the typical diurnal temperature and radiation variations over land when the sky is clear.
Figure 4-8. Clear Sky Diurnal Temperature and Radiation Variations Over Land
Warming and cooling of the Earth depend on an imbalance between solar and terrestrial radiation. The Earth receives heat during the day through incoming solar radiation.
It loses heat to space by outgoing terrestrial radiation both day and night.
Shortly after sunrise, incoming solar radiation received at the Earth’s surface (insolation) becomes greater than outgoing terrestrial radiation and the Earth’s surface warms. Peak isolation occurs around noon, but maximum surface air temperature usually occurs during the midafternoon. This lag is necessary for the air near the ground to heat up due to conduction and convection with the surface. The Earth begins to cool once the rate of outgoing terrestrial radiation exceeds the rate of insolation.
4-8 At night, insolation is absent, but outgoing terrestrial radiation continues and the Earth’s surface continues to cool. Cooling continues until shortly after sunrise, when incoming solar radiation once again exceeds outgoing terrestrial radiation. Minimum surface air temperature usually occurs shortly after sunrise.
The magnitude of diurnal temperature variation is primarily influenced by surface type, latitude, sky cover (e.g., clouds or pollutants), water vapor content of the air, and wind speed. Temperature variation is maximized over land, at low latitudes, with a clear sky, dry air, and light wind. Conversely, temperature variation is minimized over water, at high latitudes, with a cloudy sky, moist air, and strong wind.
5-1 CHAPTER 5. ATMOSPHERIC PRESSURE AND ALTIMETRY
5.1 Introduction. Atmospheric pressure is one of the most basic variables used to describe the state of the atmosphere and is commonly reported in weather observations. Unlike temperature and relative humidity, changes in atmospheric pressure are not as readily sensed by people. However, variations of pressure across the Earth are associated with pressure centers (either high pressure centers or low pressure centers) that cause the wind to blow and can bring important weather changes. Density, which is directly related to pressure, is a property of the atmosphere which can be used by pilots to help determine how their aircraft will perform at various altitudes.
This chapter discusses atmospheric pressure, how it is measured, and how it varies across the Earth. This chapter also covers the altimeter, which is a pressure sensor used by pilots to determine altitude. Finally, density will be discussed, along with its relationship to density altitude.
5.2 Atmospheric Pressure. The atoms and molecules that make up the various layers in the atmosphere are always moving in random directions. Despite their tiny size, when they strike a surface they exert pressure.
Each molecule is too small to feel and only exerts a tiny bit of pressure. However, when we add up all the pressures from the large number of molecules that strike a surface each moment, the total pressure is considerable. This is air pressure. As the density of the air increases, then the number of strikes per unit of time and area also increases.
Since molecules move in all directions, they even exert air pressure upwards as they smash into objects from underneath. Air pressure is exerted in all directions.
Atmospheric pressure is the force per unit area exerted by the weight of the atmosphere.
Since air is not solid, we cannot weigh it with conventional scales. Yet, three centuries ago, Evangelista Torricelli proved he could weigh the atmosphere by balancing it against a column of mercury. He actually measured pressure converting it directly to weight.
Figure 5-1. Air Has Weight
Air is composed of matter and, thus, has weight due to the pull of Earth’s gravity.
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5.2.1 Barometer. The instrument Torricelli designed to measure pressure was called a barometer. The aneroid barometer is the type mostly commonly used by meteorologists and the aviation community.
Essential features of an aneroid barometer (see Figure 5-2) are a flexible metal cell and the registering mechanism. Air is taken out of the cell to create a partial vacuum. The cell contracts or expands as pressure changes. One end of the cell is fixed, while the other end moves the registering mechanism. The coupling mechanism magnifies the movement of the cell driving an indicator hand along a scale graduated in pressure unit.
Figure 5-2. Aneroid Barometer
5.2.2 Atmospheric Pressure Units. Atmospheric pressure is expressed in many ways throughout the world (see Table 5-1). Meteorologists worldwide have long measured atmospheric pressure in millibars (mb or mbar), which denote pressure as a force per square centimeter. However, after the introduction of the International System of Units (SI) in 1960, the hectopascal (hPa) was adopted by most countries and is used in the Aviation Routine Weather Report (METAR)/Aviation Selected Special Weather Report (SPECI) code first developed in 1968. Many meteorologists prefer to use the term they learned during their education and work experience. Therefore, some continue to use the term millibars, while others use hectopascals (which are equivalent). The unit inch of mercury (inHg or Hg) is still used in the United States for altimetry.
5-3 Table 5-1. Units of Pressure Units of Pressure Standard Atmosphere Value at Sea Level Common Use Hectopascals (hPa) 1013.2 hPa METAR/SPECI Millibars (mb or mbar) 1013.2 mb U.S. Weather Charts Inches of mercury (inHg or Hg) 29.92 inHg U.S. Aviation Pounds per square inch (psi) 14.7 psi U.S. Engineering
5.2.3 Station Pressure. The pressure measured at an airport is called station pressure, or the actual pressure at field elevation. Pressure is lower at higher altitudes. Therefore, airports with higher field elevations usually have lower pressure than airports with lower field elevations. For instance, station pressure at Denver is less than at New Orleans (see Figure 5-3).
Figure 5-3. Station Pressure
The next few subparagraphs will examine some factors that influence pressure.
5.2.4 Pressure Variation. Atmospheric pressure varies with altitude and the temperature of the air, as well as with other minor influences, such as water vapor.
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5.2.4.1 Pressure Changes with Altitude. As we move upward through the atmosphere, the weight of the air above us decreases. If we carry a barometer with us, we can measure a decrease in pressure as the weight of the air above us decreases. Figure 5-4 shows the pressure decrease with height in the standard atmosphere.
Figure 5-4. Air Pressure in the Standard Atmosphere
These standard altitudes are based on standard temperatures. In the real atmosphere, temperatures are seldom standard, so we will explore temperature effects on pressure in the following subparagraphs.
5.2.4.2 Temperature Effects on Pressure. Like most substances, air expands as it becomes warmer and contracts as it cools. Figure 5-5 shows three columns of air: one colder than standard, one with standard temperature, and one warmer than standard. Pressure is equal at the bottom and top of each column. Vertical expansion of the warm column has made it taller than the