9. Thermal Balance of Plants and Plant Communities

9.1 Energy Balance of the Atmospheric Boundary Layer

The energy balance of plants is closely connected to processes in the atmosphere, particularly to its radiation balance and to physical transport processes, which are typically studied in meteorology (see textbooks by Lutgens et al. (2013), Stull (1988) and Wallace and Hobbs (2006)). However, the absorption of solar energy used for metabolic processes, as well as the conversion into heat fluxes and the link to water vapour fluxes, take place close to the ground, where climate conditions may largely deviate from the conditions in the free troposphere. Gregor Kraus (1911) was the first scientist to describe this phenomenon quantitatively on limestone sites near Würzburg, Germany, and thus founded a new discipline of micrometeorology (see textbooks by Jones (2014), Monson and Baldocchi (2014) and Oke and Christen (2015)). Thus, micrometeorology focuses on the boundary layer of the atmosphere—the zone near the Earth’s surface where the mean wind speed is reduced in comparison with the free airstream in the upper atmosphere.

The radiation balance of the atmosphere and the energy balance at the ground surface—the habitat of plants—are strongly dependent on the composition of the atmosphere and the optical properties of its constituents. While many trace gases and air pollutants (e.g. ozone, nitrous oxide, ammonia, methane) affect plant life, water vapour (Box 9.1), CO₂ and O₂ are the most important gases for a plant, highly relevant for its distribution, growth and fitness. Here, we will focus on the energy balance at ground level, linking radiation and sensible heat to water vapour and CO₂ in the atmosphere. Note: By definition, an energy balance is balanced (i.e. set to zero), while an energy budget can be out of balance (i.e. it can be positive or negative, or at zero).

Box 9.1: Water Vapour in the Atmosphere

The chemical composition of the lower atmosphere consists of the following gases: about 78% N₂, 21% O₂, 1% Ar, 0.6–4% H₂O, 0.040% (= 400 parts per million (ppm)) CO₂, and further trace and noble gases of natural and anthropogenic origin. For exact data on the composition of air, see List (1971) and the Intergovernmental Panel on Climate Change (IPCC) 5^th Assessment Report (2013). Here, we focus on water vapour as the most abundant greenhouse gas.

Water vapour pressure: The amount of water vapour that can be retained by the atmosphere, termed the saturation pressure (e _o), depends on air temperature and air pressure. The pressure of the atmosphere without water vapour, p _a, is derived from the measured pressure of the atmosphere, p, and the actual vapour pressure, e, as (p _a = p − e). At a constant air pressure, the vapour pressure rises almost exponentially with temperature.

Humid air is lighter than dry air, as determined by the relationship between the molecular weights of H₂O and the main constituent of air N₂, which is 0.64 (18/28). Thus, mixing of water vapour into dry air displaces some of the heavier constituents (when pressure is kept constant) and thus reduces the mass of this air volume. Hence, humid air in the lower atmosphere rises until water vapour condenses when cooling off, thus forming dew or fog at the ground or clouds in the atmosphere. For the same reason, condensation droplets form on the ceilings of humid rooms or on the lids of Petri dishes. This corresponds to dew formation in ecosystems.

The atmosphere is not always saturated with water vapour. The actual vapour pressure of the atmosphere (e) is mostly lower than the saturation value (e _o). There are a number of terms that are used to describe the humidity of air (Fig. 9.1).

The following definitions hold in meteorology:

• Absolute humidity: c _w = (2.17/T) e, unit: grams per cubic metre. (Note: The volume of air is temperature and pressure dependent.)

• Relative humidity: h = e/e _o, normally given as a percentage. (Note: At constant vapour pressure, this is dependent on T.)

• Water vapour saturation deficit of the atmosphere: D _a = (e _o − e), unit: kilopascals (often shortened to VPD (vapour pressure deficit)). (Note: The value is dependent on temperature, since e _o changes with T.)

• Water vapour saturation deficit between leaf and atmosphere: D _L = e _o(L) − e, where e _o(L) is the saturation vapour pressure at leaf temperature and e is the actual vapour pressure in the atmosphere; this term is important as the modifying driving force for transpiration in plants (often also abbreviated as VPD or—better—WSD (water saturation deficit)).

• Dew point temperature: T _d for e = e _o (the temperature, T, at which condensation temperature is reached and therefore condensation begins). This temperature is used for highly precise measurements of e.

• Wet bulb temperature: T _w. This temperature is measured to determine the actual vapour pressure (with a psychrometer): e = e _o(Tw) − γ(T _a − T _w), where γ (the psychrometric constant) is 66.1 Pa K⁻¹ for a ventilated thermometer at 100 kPa air pressure and 20 °C.

The radiation, as well as the energy balances at the ground, strongly depend on the optical properties of the atmosphere and its constituents, which interact via different biophysical processes (Mitchell 1989; IPCC 2013) and create rather complex interactions (Fig. 9.2a). The solar radiation (incident radiation on a horizontal surface) entering the Earth's atmosphere occurs in the short-wave range at about 6000 Kelvin (K), with maximum radiation at about a 0.6 μm wavelength (in the visible light). The mean radiant energy flux (measured in watts per square metre or joules per square metre per second) at the upper limit of the atmosphere is about 1361 W m⁻² (solar constant, solar irradiance; measured in the stratosphere). Since the Earth is a sphere, the mean radiation flux during the day, averaged over the entire surface of the Earth, is about 340 W m⁻² (±2 W m⁻²) (Wild et al. 2013). This incoming short-wave radiation (I _s; extraterrestrial global radiation) is partially reflected, partially balanced by outgoing long-wave radiation (thermal radiation, I _l), which by itself is a balance between thermal radiation of the atmosphere and thermal radiation from the stratosphere at a temperature of 255 K.

Long-wave radiation follows the Stefan–Boltzman law (Eq. 9.1):

(9.1)

where σ = 5.67 × 10⁻⁸ (W m⁻² K⁻⁴)—the Stefan–Boltzman constant—and T is the temperature in Kelvin. Without an atmosphere (with naturally occurring greenhouse gases), there would be no (downward, long-wave) back-radiation from the atmosphere. The average temperature on Earth would be about −18 °C (instead of about 16 °C), and thus no plant life would be possible on Earth.

However, atmospheric gases, particularly water vapour (H₂O) and CO₂, have the effect that part of the incoming solar radiation in the short-wave range is absorbed and reflected (Fig. 9.2b). Thus, the incoming short-wave radiation is limited to a narrow radiation window, with a maximum in the visible range (0.4–0.7 μm). In addition, ultraviolet (UV) radiation of the short-wave spectrum is absorbed particularly by ozone. The emitted, outgoing long-wave radiation (far-infrared) is strongly absorbed by H₂O vapour and CO₂, making the Earth hospitable for life. However, CO₂ and H₂O have an absorption minimum (emission window) between 8 and 14 μm in wavelength, in which the Earth’s surface can lose thermal energy to space. However, with increasing greenhouse gas concentrations in the atmosphere, long-wave radiation is increasingly trapped in the Earth’s atmosphere, and global temperatures are rising (Chap. 21).

The net radiation budget of the Earth (R _n; n = net) may be separated into a net radiation budget at the top of the atmosphere (R _nA) and a net radiation budget at the ground surface (R _nG) as expressed by the following equations:

(9.2)

where R _nA is the budget of radiation fluxes at the top of the atmosphere (Eq. 9.2). I _sA is the short-wave incoming (incident) solar radiation at the upper boundary of the atmosphere (extraterrestrial global radiation), ρ _sA is the ability of the atmosphere (clouds, gases) and the land surface to reflect incoming short-wave radiation, and I _lA is the total long-wave emission from the atmosphere (clouds, gases) and the land surface. Thus, R _nA can be calculated as 340 W m⁻² minus 100 W m⁻² minus 240 W m⁻², resulting in 0 W m⁻² at the top of the atmosphere.

The radiation budget at the ground surface may be expressed analogously. The “ground surface” is considered as the Earth’s surface, such as the soil surface or the top of a canopy. The net radiation budget at the Earth’s ground surface (R _nG) is the sum of the net short-wave radiation input to the ground and the difference between outgoing long-wave emission from the land surface and back-radiation—that is, long-wave radiation re-emitted to the ground by clouds and the atmosphere (Eq. 9.3).

(9.3)

where R _nG is the budget of the radiation fluxes at the ground surface. I _sG is the net incoming short-wave radiation at the ground level, calculated as the difference between incident short-wave solar radiation (I _sA), the short-wave radiation reflected by aerosols and clouds (ρ _sAI _sA), and the short-wave radiation absorbed in the atmosphere. This net short-wave radiation is also often just called light. I _lG is the long-wave radiation emitted from the land surface, and I _lAtoG is the long-wave radiation re-emitted to the ground by the clouds and the atmosphere, also called back-radiation. Thus, R _nG can be calculated as 161 W m⁻² minus 398 W m⁻² plus 342 W m⁻², resulting in 105 W m⁻² at the ground surface. The net radiation budget, R _nG, is measured with a radiometer with a polyethylene dome, which is also transparent to long-wave radiation.

Moreover, the energy balance at the ground surface (i.e. at the soil or canopy surface) also includes the heat fluxes into which the net radiation at the ground is dissipated, thus balancing the energy budget (Eq. 9.4).

(9.4)

where Φ_nG is the energy balance at the ground surface (soil or canopy), and R _nG is the net radiation budget at the ground surface. The sensible heat flux, H, is proportional to the specific heat capacity of air (c _p = 1012 J kg⁻¹ K⁻¹) and the temperature difference ΔT between the ground surface and the atmosphere, ρ is the density of air: (1.1884 kg m⁻³ at 20 °C and 100 kPa of air pressure). H, the upward flux of sensible heat, is dependent on the coupling of the exchange from the surface to the atmosphere. This coupling is expressed by the boundary resistance for heat transfer, r _b:

(9.5)

In Eq. 9.4, λE is the latent heat flux, whereby λ expresses the energy required for evaporation of water (2.454 MJ kg⁻¹ at 20 °C) and E is the evaporation from the soil and the transpiration of vegetation (kg m⁻² s⁻¹). Furthermore, Φ_nG also includes the soil heat flux—that is, the downward-oriented sensible heat flux, G (omitted in Table 9.1 because of its negligible magnitude at global level). Energy used in metabolism (M) is very small in comparison with all other energy balance components and thus is most often ignored (as in Table 9.1).

The annual incoming short-wave solar radiation (Fig. 9.3) is unevenly distributed over the Earth’s surface. Net radiation is highest over dry areas because there are fewer clouds, and it decreases towards the poles to about 40% and in the tropics to 70% of the radiation in the upper atmosphere. The global distribution of radiation is reflected in the global temperature and, in particular, low temperatures and frosts, and thus affects vegetation distribution, as well as plant performance (Chaps. 4 and 21).

About 342 W m⁻² of the 398 W m⁻² long-wave emissions from the Earth’s ground surface (I _lG in Table 9.1) are re-emitted to the ground by clouds and the atmosphere (I _lAtoG). This effect, also called the greenhouse effect, is due to clouds and trace gases and their effect on the outgoing long-wave radiation from the Earth. Water vapour, CO₂, methane (CH₄), nitrous oxide (N₂O) and other trace gases such as ozone and anthropogenic chlorofluorocarbons (CFCs) absorb long-wave radiation (I _lAtoG) and re-emit it towards the ground, affecting the temperature of the atmosphere of the Earth—that is, its climate (Stott et al. 2001). These trace gases are therefore called greenhouse gases. In particular, ozone and CFCs absorb exactly at the wavelength of the maximum long-wave emission by the Earth and are thus more effective than trace gases such as CO₂, CH₄ and N₂O, which decrease the outgoing radiation at the edge of the absorption spectrum (Fig. 9.2b). Further details will be discussed in Part 5. Trace gases in the Earth’s atmosphere are to a certain extent analogous to the window glass of a greenhouse. As window or greenhouse glass is more permeable to short-wave radiation than to long-wave radiation, short-wave radiation passes through the window glass without resistance and is absorbed by the atmosphere and ground surfaces inside the greenhouse, leading to heating of this volume. At the same time, long-wave radiation is emitted at the temperature of this absorbing volume, but this long-wave radiation cannot pass through the glass. It is thus “trapped” in the greenhouse. As a consequence, the temperature in the greenhouse further increases by heat conduction from the absorbing surfaces, since energy enters (as short-wave radiation) but does not exit (as long-wave radiation). However, the greenhouse effect in a greenhouse is intensified in comparison with the Earth’s greenhouse gas effect, since convective transport of the heated air masses is prevented in a closed greenhouse but is possible in the environment outside confined buildings.

9.2 Microclimate Near the Ground Surface

9.2.1 Daily Changes in Temperature Near the Ground

As seen in Sect. 9.1, the microclimate at ground level—most relevant to plant life—is regulated by the same processes that are active in the atmosphere. However, larger extremes occur near the ground because of the exchange of energy with the ground, which could be bare soil or the canopy of a plant community (Fig. 9.4). The consequences of short-wave energy absorption and long-wave emission on temperature become most apparent during the daily course of temperatures of bare ground during a clear day (Walter 1960; Gates 1965).

At midnight: The long-wave radiation budget is characterised by high outgoing radiation and low incoming radiation fluxes, which substantially cool the Earth’s surface. The negative radiation budget is partially compensated by heat conduction from the soil and by low heat exchange (via convection due to wind) with the air layers near the soil surface. Formation of dew and hoar frost may partially compensate this effect owing to heat released during condensation and freezing. During a cloud-free night, temperatures near the ground may fall more than 10 K below the temperature of the atmosphere 2 m above the ground. In the soil, the temperature increases with depth, reaching a time-lagged temperature minimum well after midnight, just before sunrise. At greater depths, the mean temperatures are almost constant (Scheffer 2002). This night-time cooling can limit the distribution of plant species, particularly at sites where diel temperature differences (i.e. over 24 h) can be very large—for example, at alpine elevations or in semi-deserts and deserts.

At sunrise: Incoming solar radiation quickly compensates outgoing net long-wave radiation from the soil surface. Because of the net emission to the atmosphere at night (see above) and the heat transfer into the soil, soil temperatures initially still decrease with depth during the early morning before they rise quickly at shallow depths, driven by air temperature changes, while they approach the mean temperatures of the season at greater soil depths.

At midday: Temperatures at the ground may rise up to 20 K above air temperatures measured 2 m above the ground. Heat transfer is via convection (heating of the atmosphere and development of turbulence) and heat conductance into the soil. The temperature decreases with soil depth and still shows a minimum dependent on the previous night, before reaching the mean temperature of the season at greater depth. These potentially very high temperatures during the day may impair plant regeneration at open sites—for example, after a clear-cut, in a natural forest gap or right after sowing at arable sites.

At sunset: The soil surface temperature decreases with the decreasing incoming radiation, but the temperature within the soil still rises because of the heatwave reaching the deeper soil with a time lag.

9.2.2 Modification of Environmental Radiation and Temperature by Abiotic Factors

Daily variations of temperature at the soil surface and average annual temperatures are affected by the soil conditions and site exposure (Chap. 18).

Heat conductivity and heat storage in soil: Dry soils have low heat conductivity and warm substantially faster at the surface during the day, because of incoming short-wave solar radiation, and cool greatly during the night, as the incoming (re-emitted) long-wave radiation is mainly balanced by the outgoing long-wave radiation from the soil. Moreover, in dry soils, heat exchange is limited to a small volume. On limestone near Würzburg, Germany, maximum surface temperatures of 60 °C have been measured (Kraus 1911). In contrast, wet soils have high heat conductivity and heat capacity, and thus lower daily amplitudes of temperature. Because of the higher evaporation and thus larger evaporative cooling effects on wet soils, their annual average temperatures are lower than those of dry soils.

Optical characteristics of surfaces: Absorption of incoming radiation is increased on black surfaces of rock or humus; reflection is increased on white mineral soils and with snow cover. With black organic cover, soil surface temperatures of more than 50 °C have been observed even in alpine climates. On organic soil exposed to the sun, the temperature tolerance of seedlings is often exceeded (e.g. beech seedlings are damaged at the soil surface and collapse).

Exposure and slopes: In the northern hemisphere, north slopes receive less radiation than south slopes—that is, they do not warm up as much at midday (Fig. 9.5a). South slopes receive radiation maxima according to the latitude of the site and the slope inclination. Slopes cool less during the night than the bottom of the valley, because cold air is heavier than warm air and flows downhill into the valley. The radiation budget of the slope is also more favourable during the night than in the valley, as re-radiation inputs from neighbouring slopes positively influence the radiation budget. Thus, vineyards at their northern growth limit are found on south-facing slopes, not in the valley. The climatic differences between north- and south-facing slopes in the northern and southern temperate climates may be large enough to cause differences in the local vegetation (Fig. 9.5b).

Depth and moisture of soil profile: Temperatures are much more constant within the soil than at the soil surface (see above). Depending on soil moisture, which determines the soil heat capacity, there is a clear seasonal shift of changes in soil temperature in comparison with air temperatures, with soils being generally cooler than the air during spring and early summer but warmer in autumn and early winter.

9.2.3 Modification of the Radiation Budget and Temperature by Biotic Factors

The radiation climate and temperature at the soil surface are highly affected by the vegetation cover.

Radiation reflection: The vegetation type determines its reflection. The reflection coefficient varies between 20% (e.g. for birch forests) and 12% for conifers. The low reflection of conifer forests affects the climate at the limit of boreal forests; more radiant energy is absorbed and thus more energy is kept as heat in the stand than in forest-free areas. The occurrence of evergreen conifers at high latitudes may thus accelerate a shift of the conifer treeline to the north. However, the effect is compensated by the shading of the forest floor by the evergreen canopy, which delays the snowmelt in spring and thus shortens the growing season.

Radiation absorption: The radiation budget changes with the leaf area and leaf orientation. The leaf area of vegetation cover is quantified by the leaf area index (LAI), which is defined as the sum of all leaf surfaces (one-sided leaf areas) per unit of soil area. The largest part of the incoming radiation is absorbed in the vegetation canopy, linearly scaling with LAI, modulated by the leaf arrangement in the canopy (Fig. 9.6a). Thus, light attenuation along a canopy profile in a forest follows patterns very similar to those in a cropland (see below). This can also be nicely observed in winter when snow around vegetation melts much faster than continuous snow cover without vegetation (Fig. 9.7).

With the idealised assumption of a statistically random distribution of leaves and leaf orientation in the canopy, the decrease in radiation (extinction or attenuation) follows the Lambert–Beer law of absorption:

(9.6)

where I is the radiation flux at a defined height in the vegetation, I ₀ is the incoming radiation above the canopy, LAI is the leaf area index above the measuring point and k is the extinction coefficient. k varies between 0.45 for vertical leaves (e.g. grasses) and close to 1.0 for horizontal leaves (e.g. clover). Thus, vegetation is able to increase its leaf area only until not enough light is available to maintain a positive carbon balance (Chap. 12), determined by incoming radiation (incident light, I ₀) and the angle of the leaves (modulating k). Several species (particularly legumes) are able to move their leaves. Thus, leaf inclination is no longer random, but the LAI is actively regulated. Moreover, a clumped distribution of needles in conifers results in deeper penetration of sunlight (less attenuation of light) and thus in an increased LAI in conifer stands in comparison with grassland swards (Fig. 9.6). In contrast, the extinction of incoming diffuse near-infrared radiation (short-wave) in the canopy is less efficient (i.e. less complete) in comparison with visible light (Fig. 9.6b; Jones 2014). Near-infrared radiation penetrates much deeper into the canopy than photosynthetically active radiation, because leaves are relatively transparent in the near-infrared.

Microclimate within the vegetation: With changes of radiation along the canopy profile (Fig. 9.8; Monteith and Unsworth 2013), not only temperature but also all other climate variables change, particularly the vapour pressure saturation deficit and wind speed (i.e. turbulence). Since plants not only shape but also react to their environment, feedback mechanisms exist (Chap. 13). This means that microclimatic profiles develop because of the LAI and the distribution of foliage (e.g. clumping in coniferous forests) within the stand. Thus, attenuation of light is less steep in coniferous forests (relative to the height of the canopy) than in cereal crops. At the same time, assimilation, as well as transpiration, react to low light (due to self-shading) and high relative humidity (due to soil evaporation) deeper in the canopy, in turn affecting the CO₂ concentration and relative humidity profiles and, over time (because of growth), feedback on radiation profiles. As a result, daily variations of climate variables are less extreme inside a canopy than over a vegetation-free soil surface. For example, deep in a forest, daily fluctuations in the soil temperature are small.

Thus, the gradient of climate variables within the stand depends strongly on the physiological activity of the plants (and the soil organisms) but, in addition, the canopy structure determines the roughness of its surface and in turn its interactions with the overlying atmosphere, particularly in terms of wind speed. A surface is called very rough when it strongly decreases the wind speed of an air mass passing above this surface because of its high friction. For example, tall forests or cities exhibit higher roughness (due to their higher friction) than low crop fields or bare soil, which exhibit considerably lower roughness and are considered smoother. Consequently, the exchange between the vegetation and the atmosphere, the so-called coupling, is more direct at high roughness than at low roughness, which means that the higher the roughness of a vegetation canopy is, the more deeply the wind (but also trace gases or pollutants) can penetrate into it. Thus, wind profiles, but also profiles of other climatic variables, are affected. This can be nicely seen in deciduous stands in spring, when no new foliage is present yet (Fig. 9.9).

In addition, canopy profiles vary with time on both daily and seasonal time scales. During the night, the gradients of temperature, humidity and CO₂ concentration are reversed, reflecting different physiological processes at work (e.g. respiration instead of photosynthesis), as well as differences in the coupling of the vegetation to the overlying atmosphere (Fig. 9.8). Overall, differences in microclimate within canopies and among vegetation covers depend both on their physiological activity and on the roughness of their surfaces, and thus on the coupling to the atmosphere (Chap. 18).

In deciduous temperate forests, the intercepted solar radiation is seasonally highly variable (Fig. 9.10a). Before bud burst, 40% of the incident short-wave radiation reaches the soil (transmission; I _sG). The outgoing long-wave radiation during the night is lower than that of an area without any vegetation because of the radiation shielding of canopy branches. Thus, the organic layer at the ground warms up very quickly, and spring geophytes (Fig. 9.10b) (Anemone, Corydalis, Primula and others) develop faster than the leaves of the trees. With the emergence of leaves, >90% of the energy is intercepted above-ground in the canopy (transmission 9%). Thus, the floor of the forest remains relatively cool. The light received is often not even sufficient for herbaceous plants. Therefore, shade-tolerant plants use mainly diffuse light for photosynthesis (Schulze 1972). In evergreen coniferous forests, the short seasonal “window” in which direct radiation would reach the floor, as in deciduous forests, is absent. Thus, in evergreen forests, geophytes have little opportunity to grow. Shade in the forest also delays melting of snow on the forest floor and thus delays the beginning of the growing season, with negative effects on the water balance (Chap. 10). On the other hand, this slow warming of the forest floor in coniferous forests in spring, owing to the closed evergreen canopy, is linked at the same time to a fast heating up of the canopy, owing to the low albedo of the dark surface.

9.3 Energy Balance of Leaves

The energy balance of a leaf is discussed in this section as an example, representing any organ or surface of a plant. The thermodynamic laws applicable to a leaf are basically the same as those described for the atmosphere and ground surface. However, additional parameters need to be considered for a leaf—namely, the exchange resulting from reflection of short- and long-wave radiation on the upper and lower sides of the leaf, as well as reflection and re-radiation from the soil to the leaf (Fig. 9.11).

Since I _sG represents the net radiation flux to a horizontally oriented surface, but leaves have different angles, the leaf orientation has to be taken into account when describing the short-wave incoming radiation at leaf level I _L (Eq. 9.7):

(9.7)

where β is the angle between solar rays and the object—here, a leaf (note: This equation also describes the energy gain on slopes; Fig. 9.5).

The energy balance of a leaf, Φ_nleaf, can then be estimated, taking into account the net radiation budget at leaf height, R _nL (defining “ground surface” as the leaf surface; Eq. 9.3), as well as sensible and latent heat fluxes (H and λE, respectively), energy used in metabolism (M) and the change in heat storage in the leaf (S, often ignored; Eq. 9.8).

(9.8)

The symbols and indexes are explained in Eqs. 9.1–9.3. The sensible heat flux for a leaf is:

(9.9)

Here, the boundary layer resistance, r _b, is proportional to the length of the leaf (l) and inversely proportional to the wind speed (u):

(9.10)

Ignoring the boundary layer of the leaf and assuming the leaf is well coupled to the atmosphere (Chap. 16, Sect. 16.​1), the latent heat flux is proportional to the vapour pressure deficit between the leaf and the atmosphere (D _L) and to stomatal conductance (g _s):

(9.11)

and, in addition, g _s is dependent on D _L.

However, the leaf energy balance also shows that saving water (by reducing transpiration) leads to an increased heat flux, depending on the temperature difference between the exchanging leaf surface and the surrounding air. This in turn will lead to increased leaf temperatures and might impose temperature stress (Sect. 9.4 and Chap. 4).

9.4 Acclimation and Adaptation to Temperature Extremes

The temperature limit of plant resistance to temperature extremes is generally beyond the maximum and minimum temperatures as measured in their organs under natural conditions (Chap. 4). This difference between “normal” conditions and the limits of resistance is important, as temperatures depend on incident radiation and may fluctuate rather quickly. Temperature ranges for metabolism and tolerance limits should not be exceeded, even for a short time, to avoid damage. Seedlings are particularly vulnerable to extreme temperatures, as are organs in direct contact with their surrounding boundary layer. The sequence of life forms according to Raunkiaer (trees, shrubs, herbaceous plants, annuals and geophytes; Chap. 20), with dominance of trees in the tropical regions but herbaceous plants in the cold regions of the Earth, can nevertheless not be interpreted as adaptation to temperature. There are many additional factors promoting or suppressing the growth of trees. The coldest place on Earth with vegetation (−70 °C: Oimikon, in eastern Siberia) is dominated by extensive forests of Siberian larch, and in hot climates, trees grow if water is available (e.g. palm trees in an oasis). At well-drained sites in Siberia (Chertskii), the forests extend almost up to the Arctic Ocean. The tree limit in alpine and boreal regions is caused by other factors (growing season length, frost drought, anaerobic conditions in the bogs of the tundra, fire).

Often, habitat limitation for plants is a consequence of temperature and water conditions. One example of such distribution limits is the deciduous beech (Fagus sylvatica). At the eastern borders, beech distribution is limited by the resistance of buds to the winter cold. The northern and western limits are caused by late frosts, which may damage young leaves, while the southern limit is due to drought. Moreover, “death from cold” or “death from heat” might often not be distinguishable from “death from drought” under field conditions (Chaps. 6 and 10). Nevertheless, there are many possible means of adaptation. The following examples may represent also many other species and situations.

9.4.1 Acclimation and Adaptation to High Temperatures

High temperatures combined with drought lead to many adaptations at the level of the leaf, ranging from molecular reactions (Chap. 4) to morphological changes at the levels of both acclimation and adaptation:

• Changes of leaf surface (wax, hair): The genus Encelia comprises several species characterised by variable densities of epidermal hairs (Fig. 9.12) (Jones 2014). These hairs reflect the incoming energy to different degrees (Ehleringer 1980), thus increasing survival, although at a large carbon cost (up to 70% of the annual C uptake).

• Change in leaf size: In habitats with high solar radiation, species often build smaller leaves (microphylly) (Fig. 9.13). In response to radiation stress, the same plant individual may produce small sun leaves and large shade leaves (Smith 1978). A famous exception to this rule, and therefore a clear example for adaptation, is Welwitschia mirabilis of the Namib Desert. It forms only two very large leaves, which may be up to 3 m long and 1 m wide, despite the high incoming radiation. Here, the leaf temperature is regulated by convection and long-wave radiation from the cooler soil (Schulze et al. 1980).

• Changes in leaf angle: For some plant species, particularly in the Fabaceae family, movement of the leaf plays an important role; the leaf blade moves either towards the sun to obtain maximum photon flux, or away from the incoming radiation to minimise absorption (Fig. 9.14). Classical examples are the cow pea (Vigna unguiculata) and the purple bush bean (Macroptilium purpureum). They turn their leaves to the sun or away from it, depending on the water supply. In extreme cases, they turn their leaves so that the leaf does not cast a shadow (Shakel and Hall 1979). This leaf movement changes the leaf temperature and thus the vapour pressure gradient between the leaf and the surrounding atmosphere. This in turn affects transpiration (Eq. 9.11) and thus the water relations of the plant. However, even without leaf movement, the leaf angle is an important factor in the energy balance of vegetation (Chap. 3). The hanging leaves of Eucalyptus are well known, resulting in eucalypt forests hardly creating any shadow.

• Changes of transpiration. Citrullus colocynthis is a pumpkin, which grows near the ground in desert regions of North Africa (Lange 1959). The intact leaf transpires heavily and may have a temperature of about 40 °C when the surrounding temperature of the air layer near the ground is about 53 °C (Fig. 9.15) (Larcher 2003). This is the greatest transpirational cooling that has been measured in the field. If the leaf is cut off (abscised) and thus prevented from transpiring, its temperature rises to more than 60 °C and thus exceeds the upper limit for temperature tolerance, which is around 46 °C for this plant species (Chap. 4). However, transpirational cooling can result in leaf temperatures being lower than air temperatures only at rather high air temperatures and low relative humidities, rH (e.g. at an air temperature >22 °C and an rH value of about 30%, or at 33 °C and an rH of 60%) (Jones 2014). Irrespective of the special situation of transpirational cooling below air temperature, transpiration always reduces the temperature below the level of a non-transpiring surface.

• Shading: In a free-standing tree, branches and leaves protect the stem from direct radiation. This applies particularly to subtropical dry regions, where (e.g. in Acacia reficiens) the broad canopy shades the base of the stem (Fig. 9.16). Alternatively, there is also “antenna growth” (Acacia mellifera), where shoots extend high up, as far as possible, out of the hot air layer near the ground. However, in this case, the hypocotyl still requires protection from radiation. Even in a temperate forest, gaps exist where high temperatures may occur around stems, causing damage (e.g. sunburn on beech stems in the leafless state). Thus, while having leaves or needles is a clear benefit in terms of shading (i.e. lowering temperatures within a stand), there is also a drawback of shading (i.e. preventing radiation from penetrating deep into a canopy and warming the ground, which is particularly disadvantageous in boreal and alpine climates) (Fig. 9.10).