Stable Isotopes in plant science
Isotopes are atoms of the same element that have different numbers of neutrons. Differences in the number of neutrons among the various isotopes of an element mean that the various isotopes have different masses. The superscript number to the left of an element designation is called the mass number and is the sum of the number of protons and neutrons in the isotope. For example, all isotopes of oxygen have 8 protons; however, an oxygen atom with a mass of 18 (denoted 18O) has 2 more neutrons than an oxygen atom with a mass of 16 (16O). Isotopes can be divided into radioactive and non-radioactive. The former disintegrate spontaneously over time to form other isotopes, whereas the latter do not appear to decay to other isotopes on geologic time scales, and thus are also known as stable isotopes. The most abundant elements in the biosphere are carbon (C), hydrogen (H), oxygen (O) and nitrogen (N), being 13C, 2H (D), 18O and 15N the stable isotopes of greater interest in plant physiology. The stable isotope composition of a given sample is determined by mass spectrometry, an is usually expressed in differential (d) notation:
where d stands for the isotopic composition, in parts per mil (‰),as referred to an standard (see Table I). R denotes the ratio of the heavy to light isotope (e.g., 13C/12C), and Rsample and Rstandard are the ratios in the sample and standard, respectively. A positive d value means that the isotopic ratio of the sample is higher than that of the standard; a negative d value means that the isotopic ratio of the sample is lower than that of the standard. For example, a d13C value of -28‰ means that the 13C/12C of the sample is 28 parts-per-thousand lower than the 13C/12C of the standard (Pee-Dee Belemnite limestone).
Table I Standards, notation, abundance, typical range in plants and mean analytical error of the stable isotopes most commonly used in plant physiology. Adapted from Mateo et al. (2004). Original data from Barbour et al. (2001), Ehleringer and Rundel (1988), (Epstein et al. 1977), Farquhar et al. (1989) and (Handley and Raven 1992).
a Abundances of the heavier isotope against the total pool of the element
b PDB, Pee-Dee Belemnite (limestone): already used up, replaced by secondary standards
c SMOW, Standard Mean Ocean Water
d Overall analytical precision (standard error: sample preparation + internal error of mass spectrometer)
Isotopic composition of a given element varies considerably between the different pools of the biosphere. This phenomenon is called isotopic fractionation, and is determined by isotope effects occurring during the cleavage or formation of atomic bonds, as well as during other processes affected by atomic mass (e.g. diffusion). Thus, some substances are enriched in the heavier isotope, while others become depleted (i.e. lighter). We can distinguish two kinds of isotopic effects: kinetic ad thermodynamic. The former are due to differences between isotopes in the rate of a given reaction and the latter reflect divergences in the equilibrium constants of the reaction. Kinetic isotope effects of a chain of reactions are generally non-additive, whereas thermodynamic effects are additive. Since isotope effects have values usually very close to unity, they are often expressed in terms of isotopic discrimination (D), defined as its deviation from unity:
where a is the isotope effect associated with the reaction, and dr and dp stand for the isotopic composition of reactives and products, respectively. Moreover of being a more intuitive expression of the consequences of a given process, it allows an easier comparison of the results obtained by different researchers.
The range of variation of d13C in nature is greater than 100‰ (Ehleringer and Rundel 1988). However, the most usual values in the geosphere and biosphere remains between –40‰ and 0‰ (see Fig. I). Average d13C of atmospheric CO2 is -8‰, although this value is becoming more negative year by year (ca. 0.02-0.03‰/year) due to the synergic effect of deforestation and the use of fossil fuels (Keeling et al. 1979). Moreover, there is an important seasonal variation of atmospheric d13C, which is determined by the vegetation cover, being far clearer in the Northern than in the Southern hemisphere (Fig. II). Such divergence is caused by the fact that, in Northern hemisphere, vegetation is concentrated in medium to high latitudes, and thus exposed to seasonal cycles of activity, whereas Southern hemisphere vegetation is dominated by the equatorial rainforest. Atmospheric d13C variations (seasonal or interanual) should be considered when comparing species differing in their growing cycles, as well as distant samples, either in time (at least over 4-5 years) or space (specially latitude). The vicinity (at the regional or local scale) of important sources of fossil fuel contamination should be also considered. For example, whereas in 1996 the average d13C in the Malta island was –8.0‰, the same value for Hungary was –8.3‰ (according to CU-INSTAAR and NOAA-CMDL).
The two stable carbon isotopes (13C and 12C) occur in the molar ratio of 1:99 in the atmosphere. However, plants with C3 photosynthetic pathway generally contain proportionally less 13C than does the air (Farquhar et al. 1989). Indeed, present d13C in air is about -8‰, whereas typical C3 leaf composition becomes -29‰. In many plant physiological studies, carbon isotope discrimination (D13C) is calculated as follows (Farquhar et al. 1989):
where d13Cair and d13Cplant refer to air and plant composition, respectively. This parameter reflects the amount in which the heavier isotope 13C is discriminated respect the lighter 12C during the physical and chemical processes involved in the synthesis of plant organic matter (Farquhar et al. 1989). The main factors determining D13C in C3 plants are diffusion in the air (including the boundary layer and the stomata) and carbon fixation by the carboxylating enzyme ribulose bisphosphate carboxylase (RuBisCO). Among the various models developed to describe isotopic discrimination in photosynthesis, the most extensively used is that of (Farquhar et al. 1982):
where Ca and Ci stand for ambient and intercellular partial pressures of CO2, respectively, a is the discrimination due to diffusion in air and b is the discrimination due to carboxylation. Therefore, D13C values in C3 plants depend on the relative contribution of these two steps, determined by the relationship between stomatal conductance and photosynthesis. When stomata are open (fig. IIIa), CO2 diffuses easily into the intercellular space, and Ci is closer to Ca; thus, D13C approaches the value of b (about 30‰). In other words, RuBisCO is not limited by CO2 and thus discrimination takes place mostly during the carboxylation step. In contrast, when stomatal conductance is reduced (fig. IIIb), CO2 flux is limited and Ci is significantly lower than Ca. Therefore, photosynthesis is strongly limited by stomatal conductance, and D13C becomes closer to a, the value of the discrimination during CO2 diffusion in air, (about 4.4‰).
There are several environmental factors that can modify the isotopic composition of plant tissues through their influence on either leaf conductance or photosynthetic rate, or both parameters simultaneously. Changes in irradiance levels, CO2 concentration, and plant water status (often derived from human activities) are clearly reflected in d13C variations. However, first systematic studies about plant d13C variability were focused on the identification of the recently discovered C4 and CAM photosynthetic pathways (Ehleringer and Vogel 1993; Troughton 1979). Indeed, due to the strong correlation between photosynthetic rate and stomatal conductance, it was first considered that Ci of plants would be constant, depending only on enzymatic and/or anatomical differences between different kinds of metabolism (Troughton 1979). Later, it was recognised that there is considerable variation in Ci due to environmental factors, thus reflected in d13C of plants. Since then, much effort has been taken to understand the relationship between isotopic composition and environmental variation.
Light intensity has been suggested to have a positive relationship with d13C as, in many studies, d13C of leaf tissue mimics irradiance gradients. This is the case, for example, when the vertical profile of d13C within a forest canopy is measured (Berry et al. 1997; Francey et al. 1985; Vogel 1978). However, interpretation of these results has been controversial as it is often difficult to discriminate direct light effects from those caused by differences in either CO2 concentration, vapour pressure or air d13C throughout the canopy. Vogel (1978) attributed this “canopy effect” to a recycling of soil CO2, which had lower (about 12‰) d13C than atmospheric CO2. However, in those studies where simultaneous measurement of air d13C were performed (Berry et al. 1997; Francey et al. 1985), plant d13C was significantly correlated with sampling height, without a corresponding change in air d13C. Therefore, most of the decrease in leaf d13C with canopy depth is likely to be related with stomatal and photosynthetic effects. This does not fully discard CO2 and vapour pressure gradients as potential causes for the observed d13C variability within canopies (Broadmeadow et al. 1993).
Other works have shown strong evidence of a direct effect of light on d13C, at least when this factor is limiting (Broadmeadow et al. 1993; Yakir and Israeli 1995; Zimmerman and Ehleringer 1990). Taking advantage of the negative relationship between leaf oxygen isotope composition (d18O) and transpiration rates, (Yakir and Israeli 1995) demonstrated that the cause underlying the relationship between d13C of banana leaves and irradiance was an increase in photosynthetic capacity, without appreciable variations in stomatal conductance. Other authors arrived to similar conclusions combining gas-exchange measurements with d13C analyses (Broadmeadow et al. 1993). In contrast, (Zimmerman and Ehleringer 1990) sustained that the origin of d13C increase in orchid leaves at high irradiance was due to declining stomatal conductance. This conclusion should be taken carefully, as it was only based on the small differences found in leaf nitrogen content (concentration?), assuming that this would imply little change in potential photosynthetic capacity (Lambers et al. 1998).
Early in the 60s, (Park and Epstein 1960) grew tomato plants at two levels of CO2 and showed that plants grown at the higher level had a more negative d13C than plants grown at the lower concentration. Since then, similar results have been reported for several C3 herbs (Beerling and Woodward 1995; Polley et al. 1993) and trees (Beerling 1997; Picon et al. 1997) over a wide range of CO2 concentrations. In C3 plants, CO2 is usually limiting photosynthesis and, thus, an increase in CO2 results in greater photosynthetic rates. On the other hand, plants take advantage of the increased CO2 availability to augment water use efficiency (i.e. the ratio between net assimilation and water transpired) by closing stomata. Experimental results indicate that this reduction in stomatal conductance does not limit photosynthesis, thus d13C values become more negative as CO2 concentration increases (Beerling and Woodward 1995; Polley et al. 1993). Therefore CO2 and the above-mentioned light gradients may have additive effects within closed canopies, both contributing to leaf d13C decrease with depth (Berry et al. 1997; Broadmeadow et al. 1993).
Plants typically react against a decrease in water availability through stomata closure and, although carboxylation rates may also decline under water shortage, leaf conductance is usually affected to a larger extent, originating a reduction in Ci and a concomitant increase in d13C (Farquhar et al., 1989) (Scheidegger et al. 2000). Many studies under growth-chamber and field conditions have shown that plants developed under water stress (stress induced by low soil water content) produced leaves with higher d13C (see references in (Ferrio et al. 2003; Griffiths and Parry 2002; Mateo et al. 2004; Warren et al. 2001). On the other hand, it is leaf water availability the factor that ultimately influences d13C and this availability depends not only on the water input from the soil, but also on its physical structure as well as on the hydraulic resistance along the plant xylem (Masle and Farquhar 1988; Warren and Adams 2000). Moreover, the rate of evaporation from the leaf also determines stomatal responses that subsequently affect d13C. Indeed, an increase in the leaf-to-air vapour pressure gradient (VPG, which is considered the driving force for transpiration) will also cause a reduction in Ci, leading to higher d13C values (Barbour and Farquhar 2000; Ehleringer 1990). In accordance with these assumptions, it should be expected to find significant relationships between d13C and environmental parameters related with water availability or VPG, such as precipitation, relative humidity or potential evapotranspiration. (Stewart et al. 1995), for example, analysed d13C from 12 plant communities along a rainfall gradient, finding significant negative correlations between annual rainfall and community-averaged d13C. Similar relationships with water availability have been reported over smaller scales, reflecting microenvironmental changes (Ehleringer and Cooper 1988; Peñuelas et al. 1999). On the other hand, the expected positive correlation between VPG and d13C along climate gradients has also been reported elsewhere (Sparks and Ehleringer 1997; Williams and Ehleringer 1996).
It is important to note here, however, that in most cases the relationship between d13C and plant water availability is not linear, showing a saturation trend as water availability increases (Araus et al. 1997a; Korol et al. 1999; Warren et al. 2001). (Warren et al. 2001), for example, found that, over a global survey of d13C values on conifers, d13C reached an asymptotic value once there is no water deficit, namely when the ratio between precipitation and evapotranspirative demand equalled unity. The reason for that general trend is that the main factor relating d13C with water inputs is stomatal conductance, which is expected to reach its maximum in non-stressed plants. Under optimum water availability, no further increments in stomatal conductance, and thus on d13C, would be expected (Lambers et al. 1998).
The interpretation of oxygen isotope composition (d18O) variability in plants is not as direct as that of d13C, as it involves a combination of abiotic and biotic fractionation processes. The study of oxygen isotopes has been traditionally linked to that of the other element of water, hydrogen. In 1932, (Urey et al. 1932) discovered a heavier form of hydrogen: Deuterium (2H). Using measures of water density, they found considerable variations among samples of different origins. Years after the discovery of Urey, (Dansgaard 1964) began a systematic analysis of 18O and 2H in marine and continental waters, finding that the first were more enriched in the heavier isotopes than the latter. On the other hand, the relationship between d18O and d2H in the world’s fresh waters follows a predictable linear relationship, referred to as the “meteoric water line” (Craig 1961).
Craig’s meteoric water line is a direct consequence of the fact that fresh waters of warm regions have more enriched (positive) values of hydrogen and oxygen isotopes, whereas cool regions are isotopically depleted for both elements. Such differences are mostly due to fractionation occurring during phase changes in the course of the hydrological cycle (see Fig. IV). On the one hand, light isotopes (16O and 1H) evaporate more rapidly than their heavier counterparts, and thus water vapour is isotopically depleted respect source water (e.g. ocean water). On the other, the opposite occurs during precipitation. In order for water to condense and precipitate from an air mass, the temperature must drop. As temperature decreases, the heavier isotopes are selectively precipitated through distillation. This phenomenon is referred to as the “rainout” effect (Dansgaard 1964), which describes the successive depletion as air mass travels from its source and loses heavier isotopes. This effect can be enhanced by orographic effects that may cause an abrupt cooling of an air mass as it rises. On the other hand, the amount of this fractionation varies according to temperature and, as a consequence, isotopic composition of precipitation is positively correlated with temperature.
Fig. IV Main fractionation steps and typical values of oxygen isotope composition (d18O) in a temperate climate. d18OSMOW, standard mean d18O in ocean water; d18OV, d18OP, d18OLW, d18O in water vapour, precipitation (either rainfall or snow) and leaf water, respectively; VPD, vapour pressure deficit; gs, stomatal conductance. Redrawn from Ferrio et al. (2005b). Original data from (IAEA/WMO 2001) and (Saurer et al. 1997b).
The source of water for most terrestial plants is soil moisture, so part of the signal in water isotopes (H and O) of plants come from the isotopic signature of precipitation. However, there are several potential fractionations before the water isotopes become fixed in plant tissues (Fig. IV). The first occurs within the soil, as evaporation affects the original isotopic signal, so the residence time and deep of soil water is important (Buhay and Edwards 1995). On the other hand, the observed discrepancy between the d18O of molecular oxygen in atmosphere (of photosynthetic origin) and marine water lead to suspect about some fractionation at the plant level. (Gonfiantini et al. 1965) proved for the first time that the isotopic composition of leaf water was enriched in heavy isotopes during transpiration. When plant roots take soil water there is no fractionation, so the critical site of fractionation is the leaf. Summarising, the d18O of plant tissues reflects the variation in (1) d18O in source water, (2) evaporative enrichment of leaf water due to transpiration, and (3) biochemical fractionation during the synthesis of organic matter (Farquhar and Lloyd 1993; Yakir 1992). The level of enrichment of leaf water above source water (D18Oe) has been modelled as follows (Dongmann et al. 1974; Farquhar and Lloyd 1993):
where e* is the proportional depression of vapour pressure by the heavier H218O, ek is the diffusion fractionation through stomata and leaf boundary layer, ea and ei stand for vapour pressure in atmosphere and intercellular space, respectively, and D18Ov is the oxygen isotope composition of water vapour in the atmosphere (relative to source water). At constant temperature, and where source water and atmospheric vapour have the same isotopic signature, D18Oe is linearly dependent on 1- ea/ei (Barbour et al. 2001). According to this model, plants growing at higher humidity (i.e. higher ea) are expected to have lower D18Oe. Consequently, within the same environmental conditions, plants with higher stomatal conductance (i.e. lower leaf temperature, which reduces ei) are also expected to show smaller D18Oe. This models explains the isotopic enrichment of water at the site of evaporation, however, backward diffusion of the enrichment is opposed by the convection of isotopically lighter source water to the sites of evaporation, what is known as the Péclet effect (Barbour et al. 2004; Barbour and Farquhar 2000). The d18O of leaf water is, therefore, less depleted than that at the site of evaporation, so Eq. V overestimates the net effect of transpirative enrichment. Leaf water enrichment is passed on to the organic molecules formed in the leaf by exchange of oxygen atoms between carbonyl groups and water (Sternberg et al. 1986). The d18O of sucrose formed in the leaf is proportional to the signal of leaf water, although with 27‰ enrichment. However, in the case of stem cellulose, most of the enrichment signal (derived from sucrose) is further exchangeable with xylem water during the heterotrophic pathways of cellulose biosynthesis (Roden et al. 2000; Sternberg et al. 2003). This is particularly important for tree-ring studies.