Jessica Gurevitch, Samuel M. Scheiner, and Gordon A. Fox - The Ecology of Plants - 2002 / The Ecology of Plants Jessica Gurevitch, Samuel M. Scheiner, and Gordon A. Fox; 2002 / Ch 18

352 Chapter 18

vegetation map was produced. Over the next several decades, a number of naturalists developed classifications of plant communities, noting the relationships between different types of communities (such as forests and grasslands) and the climate at different latitudes and altitudes.

Our understanding of weather has depended heavily on the ability of researchers to gather large numbers of measurements over a wide geographic area. Prior to the mid-eighteenth century, no one had realized that weather moves in predictable ways across the globe— news traveled far more slowly than weather (McIlveen 1992). On October 21, 1743, Benjamin Franklin attempted to observe a lunar eclipse in Philadelphia, but was prevented from seeing it by a storm. Later, he was surprised to learn that the eclipse was visible in Boston and that the storm arrived there the following day. By contacting people living between the two cities, he was able to reconstruct the movement of the storm. It was the development of communication technology, however, that really changed the sciences of meteorology and climatology. The invention of the telegraph in 1835 (with the first message transmitted in 1844) made it possible to organize large numbers of people to observe and forecast the movement of storms. The advent of airplanes and the military technology developed in World War II further spurred the development of weather station networks and made it possible to sample weather conditions high in the atmosphere. Today, we gather information about the weather with a combination of satellites, ground-based stations, and upper-atmosphere probes. The ready accessibility of powerful computers makes it possible to analyze these enormous data sets, and both climatologists and meteorologists are actively developing more sophisticated models of climate and weather (see Box 22A). These efforts are especially important in our attempts to predict changes in weather and climate due to the effects of global warming and to understand how different factors influence those changes.

The two primary components of a region’s climate are temperature and moisture. We first turn to the longterm mean (average) temperature and precipitation and the effects of these average conditions on vegetation patterns. Variation around these means is also important, and we look at two aspects of this variation: predictable changes (e.g., it is usually warmer in the summer than in the winter) and departure from average conditions (e.g., in a particular year it may be unusually dry in the Amazon basin). The time scale of variation is also important, with temperature and precipitation varying on scales from a single day, a week, a season, a year, a decade, up to cycles that take hundreds of thousands of years to complete. Here we investigate these patterns of variation and how they shape plant communities.

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Temperature

If you sit outside on a sunny day, you can feel the sun’s radiant energy on your skin. Radiant energy from the sun, and variation in how much radiant energy reaches different parts of Earth’s surface, is the primary cause of variation in temperature, the first of the two major climatic factors. Heat is a measure of the total kinetic ener- gy—the energy of motion—of the molecules in a substance, while temperature is a measure of the average kinetic energy of those molecules; see Chapter 3. Radiant energy from the sun heats objects (such as plants or the ground) directly. This energy is primarily in the shortwave part of the electromagnetic spectrum (wavelengths < 700 nm) and includes visible light. As objects are heated by sunlight, they emit longwave radiant energy (wavelengths >700 nm, including infrared radiation) in proportion to the fourth power of temperature (see Chapter 3), becoming a secondary source of heating for nearby objects. The warming of physical objects by sunlight, and their slow emittance (i.e., the storage and later release) of that heat energy over time as longwave radiation, creates lags in temperature change, which we will explore in more detail below. The transport of that heat energy from one place to another by atmospheric and ocean currents—a process called convection—is also a key element in global climates.

The temperature at any spot on Earth’s surface is determined primarily by the amount of radiant energy it receives from the sun. That amount is determined, in turn, largely by the angle of Earth’s surface in that spot relative to the sun’s rays, as well as by atmospheric conditions. Consider a spot at sea level on the equator at noon during the vernal or autumnal equinox, when the sun is directly overhead. At the top of the atmosphere at that spot, Earth receives approximately 340 W/m2/year, or 2 langleys (ly) of energy per minute (1 ly = 1 cal/cm2) as shortwave radiation. If we consider the amount of energy striking the top of Earth’s atmosphere to be 100%, we can trace what happens to that energy (Figure 18.1). Almost a third of it returns to space as shortwave radiation, reflected by Earth’s surface (6%) and by clouds (17%) and scattered by the atmosphere (8%). Almost half (46%) is absorbed by Earth’s surface. The rest of the energy entering the atmosphere is absorbed by clouds (4%) and by the “greenhouse gases” (19%) in the atmosphere. (We will return to the absorption of energy by greenhouse gases shortly and will discuss it at greater length in Chapter 22.)

Earth’s surface emits energy in the longwave part of the spectrum, reradiating 15% of the initial total energy that reaches the upper atmosphere and transferring an additional 24% as latent heat loss (see Chapter 3) and 7% as convection to the atmosphere and clouds. The greenhouse effect results when of greenhouse gases in the

Figure 18.1

The radiant energy balance of the Earth, expressed as the percentage of the energy reaching the top of the atmosphere (340 W/m2/year, or 100%). This energy is reflected by the surface and by clouds and scattered by the atmosphere, so that about 31% returns to space. Of the remainder, 46% is absorbed by the surface and 23% by the atmosphere, warming both. Longwave radiation (infrared, shown by boldfaced numbers) is emitted by the surface, warming the atmosphere further as well as being lost to space. Convection and latent

atmosphere reabsorb longwave radiation emitted by the Earth’s surface (see Figure 22.3). Secondary heating of the Earth’s surface then occurs by reradiation from the atmosphere.

This reabsorption and reradiation is critical for life on Earth. Without the greenhouse effect Earth would be bitterly cold everywhere, all of the time. Secondary heating by longwave energy from the atmosphere accounts for two-thirds of the total radiant energy received at the surface. The total outgoing longwave radiation from Earth back to space is about 69% of the energy received initially, coming from Earth’s surface, the atmosphere, and clouds.

Incoming energy must balance outgoing energy. Temporary imbalances will result in Earth warming or cooling, but eventually a new equilibrium will be reached as Earth’s emission of longwave radiation changes according to its new temperature. Thus, the

heat exchange also transfer heat energy from the surface to the atmosphere. The total amount of energy absorbed by the surface, atmosphere, and clouds (as shortwave radiation) and the total amount that leaves the planet (as longwave radiant energy, latent heat loss, and convection) are in balance with each other. The total amount of energy entering the top of the atmosphere (as shortwave radiation) is exactly equal to the sum of the outgoing shortwave radiation (from reflection and scattering) plus the outgoing longwave radiation, which together sum to 100%. (After Schlesinger 1997.)

shortwave radiant energy absorbed by the atmosphere and Earth’s surface (46% + 23%) is balanced by the longwave energy (69%) emitted.

Short-Term Variation in Radiation and Temperature

The combination of shortwave and longwave radiation determine the ambient temperature at Earth’s surface. As you travel away from the conditions described above, of maximal solar energy input at the equator at noon, several factors act to determine the amount of incoming radiation and the subsequent temperature. Most important is the angle of the sun, which determines the amount of incoming shortwave radiation. Over the course of the day, Earth rotates on its axis, and the sun’s rays hit the surface at different angles. The solar angle of incidence is the angle that a ray of sunlight makes with a line perpendicular to the surface. The smaller the angle of incidence, the closer to 0º from the horizontal, the larger the

Less solar energy is intercepted by horizontal surface parallel to surface

of Earth

Sun’s rays

Maximum amount of solar energy is intercepted by horizontal surface parallel to surface of Earth (midday)

Subsolar point (90º from

horizontal)

0º from horizontal

Shadowed hemisphere (night)

Energy imput to horizontal surface (W/m2)

1500

1000

500

Sun’s angle from horizontal

Figure 18.2

Differences in angle and amount of solar radiation reaching the Earth’s surface. At the equinox, when neither hemisphere points toward or away from the sun, the diagrams shows the amount of solar radiation received at a single spot over the course of the day. The same differences could apply to the angles and amounts of solar radiation reaching different latitudes at different times of the year (see Figure 18.3B).

area over which a given amount of solar energy is spread. Consequently, when the incoming sunlight is at a steep angle, any given square meter of surface receives less energy than it would if the sun were at a more direct angle overhead (Figure 18.2), Also, at steeper angles, the incoming solar energy must travel a longer distance through the atmosphere, and so more energy is absorbed by the atmosphere and less reaches Earth’s surface.

Similarly, as one moves north or south of the equator, the angle of the sun, even at noon, decreases. The actual angle depends on the time of year and how far one is from the equator. Earth’s axis is tilted relative to the plane in which it revolves around the sun. Seasons are created by the differences in solar angle, and in the consequent amount of incoming solar energy, over the year-long progression of Earth around the sun (Figure 18.3A). The Northern and Southern Hemispheres are tilted toward or away from the sun at the opposite extremes of Earth’s trajectory around the sun, at the solstices in December and June. At the equinoxes, in September and March, Earth’s axis is parallel to its plane of revolution, neither pole is tilted toward or away from the sun, and day and night length are equal everywhere on Earth. When the sun is more directly overhead in the Northern or Southern Hemisphere, it is summer in that hemisphere. In the other hemisphere, steep solar angles result in decreased radiant energy input, and it is winter. This change in solar angle with the seasons is greatest at the poles and least at the equator. As a result, over the course of a year at extreme latitudes, there are large changes in radiation input. In contrast, equatorial regions have

nearly constant solar energy inputs throughout the year (Figure 18.3B). These effects combine to produce distinct patterns of temperature variation across the globe (Figure 18.4).

Atmospheric, ground, and oceanic temperatures have a feature called lag—a delay in effect—that determines some of the daily and yearly patterns of temperature variation. It is warmer in the late afternoon, even when the sun is low in the sky, than when the sun is at a comparable angle in the morning. The coldest month of the winter in the Northern Hemisphere is not December (when the sun is lowest and the days are shortest), but January; conversely, the hottest summer months are July and August, not June (when the days are longest and the sun most directly overhead). The cause of lag is the storage of heat by the ground, the ocean and other large bodies of water, and the atmosphere itself.

Imagine a place somewhere at a temperate latitude at the winter solstice. Incoming solar radiation is at a minimum. As Earth revolves around the sun and the year progresses, day length increases and the sun is more directly overhead. The ground absorbs solar energy and begins to warm. The ground has a certain specific heat—the amount of heat energy that is required to warm 1 gram of it by 1°C. Water has a higher specific heat than the ground, so the ocean takes longer to warm up than nearby land. Air has a far lower specific heat than either land or water, and warms up more readily than they do. Even as our imaginary spot on Earth passes the summer solstice and incoming radiation begins to decline, the longwave output from the now warmed ground and bodies of water nearby continue to warm the air, keeping air temperatures high.

Eventually, as solar input declines, the reverse pattern occurs: air, ground, and finally water become cooler. A high specific heat results in slow cooling, just as it causes slow warming. Coastal environments, influenced by the adjacent bodies of water, consequently cool and warm more slowly than those in the interior of conti-

Winter

0º

43º

December 66.5º solstice

90º

23.5º

Summer

nents, and thus experience more moderate temperatures. The water ameliorates very hot temperatures as cooler offshore breezes convectively reduce heat on land. In autumn, the water releases heat gradually, slowing temperature declines on land by providing warming breezes. These breezes may feel cold to a person, but they are warm relative to the ground and the air over it. One subtle implication of these lags for plants is that times of maximum or minimum solar radiation are not the times of highest or lowest temperatures, either daily or seasonally.

The magnitude of daily variation in temperature— the range from nighttime to daytime temperatures— depends in part on local atmospheric conditions, particularly the amount of moisture in the air. At night, there is no incoming shortwave radiation. We do not immediately freeze after sunset, however, because the surrounding air acts somewhat like a blanket, holding heat and releasing it as longwave radiation all night. Moist air can hold more heat than dry air, as we will see below. A place like Tokyo, adjacent to the Pacific Ocean with

typically humid air, usually has nighttime temperatures that are only a little cooler than daytime temperatures. In contrast, Tucson, Arizona, with very dry air, may have dramatic swings, with nighttime temperatures 20°C or more cooler than those during the day. As altitude increases, the density of the air decreases—the blanket is thinner. As a result, areas at high elevations, such as

mountaintops, also tend to have large swings in temperature from day to night. Plants must be adapted to tolerate the temperature ranges, as well as the average conditions, where they live. Elevation and latitude jointly affect climate in interesting ways. As indicated above, equatorial regions receive nearly constant amounts of radiation throughout the year. But there are high-ele- vation regions at the equator, such as the Ecuadorian Andes in South America and the Kenyan highlands in

Africa. These spots can experience daily temperature swings approaching those experienced between summer and winter at mid-latitudes, although obviously they also receive much more radiation during the day.

Long-Term Cycles

Besides the daily and yearly variations in solar radiation, there are also longer-term cycles. One of these is the 11year sunspot cycle. Over this cycle, the amount of solar

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radiation reaching Earth varies up to 0.1%. A 22–23-year double sunspot cycle is also known, and is related to reversals of the solar magnetic field. These cycles probably do not directly affect plant growth or local temperatures. Rather, they affect atmospheric circulation and precipitation patterns, as we will see below, resulting in indirect effects on plants due to changes in weather patterns. The linkage between the sunspot cycle and climate

Climate and Physiognomy 357

was first studied by the astronomer Andrew E. Douglass, who used the long-term record of good and bad years for plant growth recorded in tree rings. In doing so, he invented the field of dendrochronology (see Chapter 13).

At very long time scales, there are several different kinds of changes in Earth’s orbit around the sun that

^ influence climate, collectively called Croll-Milankovic

effects (Figure 18.5). The longest of these orbital cycles

(A) Ellipticity of orbit

Cyclic period = 100,000 years

Aphelion

(B) Degree of axial tilt

Figure 18.5

N

Most circular

orbit Most elliptical orbit

Perihelion

Minimum tilt

Polaris (North Star)

N

Present

Ellipticity of orbit

Degree of axial tilt

Direction of axial tilt

0.03

0.01

24.5

24.0

23.5

23.0

22.5

22.0

Figure 18.6

^ Past and predicted Croll-Milankovic cycles.

Sometimes the cycles can reinforce each other; for example, when both ellipticity and the degree of axial tilt are large. At other times they can cancel each other. The precentage of axial tilt is the amount of wobble. (After Gates 1993.)

is a 100,000-year cycle of change in the shape of Earth’s orbit around the sun, from nearly circular to a more elliptical trajectory. Currently the orbit is almost circular, so that the difference between Earth’s nearest point to the sun and its farthest is only 3.5%. When the orbit becomes more elliptical, this difference between the shortest and longest distance of Earth from the sun may be as great as 30%. Clearly, such changes in orbital shape will affect the range of seasonal variation at temperate and polar latitudes, because the amount of incoming solar radiation will change.

The next longest pattern of variation is a 41,000-year cycle in the degree of tilt of Earth’s axis from 22.1° to 24.5°; currently, the tilt is 23.5°. These changes in the axial tilt will greatly affect the range of seasonal variation, because changes in the angle of the sun will change the amount of solar energy input at any given spot in temperate and polar latitudes.

Finally, there is a 22,000-year wobble of Earth’s axis. This wobble is similar to what you would see on a top that is beginning to slow down, so that its stem no longer points straight up but is beginning to circle around. The result of this wobble is to change whether the Northern or the Southern Hemisphere is pointing toward the sun during the time of Earth’s nearest approach to the sun. Earth is now a little bit closer to the sun at the winter solstice in the Northern Hemisphere (the summer solstice in the Southern Hemisphere) than at the opposite point in its orbit. Changes in the direction of the axial tilt will affect the intensity of seasonal variation; for example,

Northern Hemisphere winters currently are milder than Southern Hemisphere winters.

All three of these cycles thus modify seasonality, and they can work together or in opposition to magnify or minimize seasonal variation (Figure 18.6). While the effects may be small, they are sufficient to cause largescale and long-term changes in Earth’s climate and may be responsible for the cycles of ice ages that Earth has experienced over the last million years (see Figure 21.8). We look at the most recent of these ice ages and its aftermath in Chapter 21.

Recently there has been extensive debate about global warming, an overall rise in global temperatures. Part of the debate stems from a genuine scientific question: is the recent warming of the climate likely to be caused by human addition of greenhouse gases to the atmosphere, or does it reflect natural cyclic changes and perhaps some random variation in the weather? Without considering the actual data, either hypothesis (or both) could be plausible. Are the warmer temperatures we are experiencing part of a long-term trend, or are they merely a short-term blip that will soon be reversed? If this warming is part of a trend, to what extent is it due to human actions? Certainly, it is very difficult to separate natural variation in weather and climate from anthropogenic changes. In any case, one may well wonder what the answers to these questions mean for the future: Is the observed warming a harbinger of an increasingly warm climate? We investigate this issue in depth in Chapter 22.