Soil Phosphate Stable Oxygen Isotopes across Rainfall and Bedrock

Article pubs.acs.org/est

Soil Phosphate Stable Oxygen Isotopes across Rainfall and Bedrock Gradients Alon Angert,*,† Tal Weiner,† Shunit Mazeh,† and Marcelo Sternberg‡ †

The Institute of Earth Sciences, The Hebrew University of Jerusalem, Israel Department of Molecular Biology & Ecology of Plants, Wise Faculty of Life Sciences, Tel Aviv University, Israel

‡

S Supporting Information *

ABSTRACT: The stable oxygen isotope compositions of soil phosphate (δ18Op) were suggested recently to be a tracer of phosphorus cycling in soils and plants. Here we present a survey of bioavailable (resin-extractable or resin-P) inorganic phosphate δ18Op across natural and experimental rainfall gradients, and across soil formed on sedimentary and igneous bedrock. In addition, we analyzed the soil HCl-extractable inorganic δ18Op, which mainly represents calcium-bound inorganic phosphate. The resin-P values were in the range 14.5−21.2‰. A similar range, 15.6−21.3‰, was found for the HCl-extractable inorganic δ18Op, with the exception of samples from a soil of igneous origin that show lower values, 8.2−10.9‰, which indicate that a large fraction of the inorganic phosphate in this soil is still in the form of a primary mineral. The available-P δ18Op values are considerably higher than the values we calculated for extracellular hydrolysis of organic phosphate, based on the known fractionation from lab experiments. However, these values are close to the values expected for enzymatic-mediated phosphate equilibration with soil−water. The possible processes that can explain this observation are (1) extracellular equilibration of the inorganic phosphate in the soil; (2) fractionations in the soil are different than the ones measured at the lab; (3) effect of fractionation during uptake; and (4) a flux of intercellular-equilibrated inorganic phosphate from the soil microbiota, which is considerably larger than the flux of hydrolyzed organic-P.

1. INTRODUCTION The possible use of stable oxygen isotopes of soil phosphate as a tracer for P cycling has gained attention lately.1−4 This use is based on the stability of the P−O bond, which under natural soil conditions is broken only by enzyme-mediated reactions.5−7 These reactions cause either kinetic isotope effects or isotopic equilibration of the PO4 oxygen with the oxygen in H2O. The equilibration is believed to be mediated by pyrophosphatase,8 which is a ubiquitous intracellular enzyme that appears in all living cells and results in temperaturedependent isotopic equilibrium with water:6,9 δ18Op = δ18Owater + (111.4 − T )/4.3

in DNA and RNA, which are the dominant forms of organic-P in cells, requires two steps. In the first step, cleavage of Pdiester results in the generation of phosphomonoester (Pmonoester), while in the second step, cleavage of the Pmonoester releases PO4 (inorganic-P or Pi). The Pi released by these two steps will inherit half its O atoms and their isotopic signature from the organic-P, while the other half of the O atoms are the result of incorporation of water O, with enzymeand substrate-dependent isotope effects.10 The isotopic composition of the resulting Pi is given by δ18OP = 0.5(δ18OP‐org ) + 0.5(δ18O W + F )

(1)

where F is the combined fractionation factor associated with both steps: F = (F1 + F2)/2. The dominating soil monophosphate-esterase in the study area is alkaline phosphatase.11 The combination of deoxyribonuclease or phosphodiesterase

where T is the temperature in degrees Celsius and δ18Op and δ18Owater represent, in standard delta notation, the oxygen isotope composition of phosphate and water. While inorganic pyrophosphatase is normally located inside the cell, it could be released to the extracellular environment via cell lysis. Causes of the major kinetic effects reported thus far include hydrolysis of P-bearing organic molecules. Hydrolysis of phosphodiester (P-diester) or phosphodiester bonds, such as © 2012 American Chemical Society

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Received: Revised: Accepted: Published: 2156

October 6, 2011 January 3, 2012 January 9, 2012 January 9, 2012 dx.doi.org/10.1021/es203551s | Environ. Sci. Technol. 2012, 46, 2156−2162

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to compare the measured values and their range to the values expected from our current understanding of the processes affecting δ18OP.

(PDase) in the first step, and alkaline phosphatase (APase) in the second step, results in the following fractionation factors: For DNA hydrolysis, F is −25(±6)‰ for the diesterase + APase pathway (F1 = −20‰, F2 = −30‰). For RNA hydrolysis, F is −5(±6)‰ for the diesterase (PDase 1) + APase pathway (F1 = +20‰, F2 = −30‰).10,12 Hydrolysis of organic-P, which is originally in the form of P-monoester (e.g., glucose 6-phosphate or phosphocholine),13 will involve only the second step, and as a result, 0.75 of the organic isotopic signature will be retained, while the rest will be incorporated from water, with a kinetic fractionation of −30 ± 8‰ for APase.12 In soils, RNase or some other nonspecific PDase can be also responsible for the first step of RNA degradation and may produce a different fractionation. The fractionation factor associated with another common soil monophosphate-esterase, acid phosphatase, is currently unknown, and the fractionation associated with 5′-nucleotidase was reported to be −10(±1)‰.12 It should be noted that the fractionations obtained in the lab from experiments with commercially available enzymes may be different than those that occur in a natural soil system. Phosphate released from organic P re-enters the inorganic pool and then can be assimilated by plants or by the soil microbial population or be sorbed to soil and form secondary phosphate minerals. Lab experiments have shown that precipitation of apatite involves only very small isotopic fractionations of ∼1‰.14 A recent study15 has found negligible fractionation between phosphate sorbed into iron oxides and the one in aqueous phase at equilibrium, and only small fractionations (∼2‰) in the early stages of phosphate sorption. Other sorbing reactions in soils might have different isotopic effect associated with them. No isotopic fractionation takes place in the dissolution of apatite, which is the main source of primary P in soils.16 Inorganic phosphate is assimilated from the soil solution into plants or soil microbial biomass. Fractionations involved in this step, if any, are unknown. Fractionation of −3‰ was observed for Escherichia coli under laboratory conditions.8 However, this effect may not represent the soil microbial population under natural conditions and can be different from that of fungi and plants. Given the above, we expected that the strong isotope effects associated with organic-P hydrolysis will produce Pi with values substantially different from those produced at equilibrium catalyzed by inorganic pyrophosphatase. In addition, dependence of the isotope effects on the substrate and enzymes involved, as well as variations in the δ18O of organic-P, can create potentially large variability between different ecosystems. Here we have determined, for the first time, the variability in δ18OP across a climatic and bedrock gradient for both resinextractable P and HCl-extractable P. The resin-extractable P (hereafter resin-P) is the phosphate that can be collected on an anion-exchange resin by shaking it in water to which a soil sample was added. This phosphate fraction is considered to be a form of labile Pi adsorbed on surfaces of crystalline compounds from which plants normally draw their supply, while the soil HCl-extractable P (hereafter HCl-P), is considered to be mostly Pi that is calcium-bound or occluded within sesquioxides.17−19 This research focused on arid and Mediterranean soils, in which P is considered a limiting factor and where the small pool of labile-P can experience fast transformations. The study aimed to survey the variations in soil δ18OP, to assess the natural variability in this parameter, and

2. METHODS 2.1. Study Sites. Soil samples were collected from five sites in Israel with natural vegetation, in which no fertilization takes place (Table 1). Sites 1−4 are located along a 245 km long Table 1. Description of Research Sites MAPa (mm·year−1)

site

location

1

300

3

central Negev Desert northern Negev Desert Judean Mountains

4

Galilee Mountains

780

5

Golan Heights

700

2

a

90

540

bedrock hard limestone hard limestone hard limestone hard limestone basalt

comments

rain manipulations rain manipulations

MAP, mean annual precipitation.

rainfall gradient, going from arid (site 1) through semiarid (site 2), and Mediterranean (site 3) to mesic Mediterranean (site 4), where the mean annual rainfall is 90, 300, 540, and 780 mm, respectively. The rainy season is between October and May. These four sites were established as part of the GLOWA Jordan River (JR) project, and their full description has been previously published.20,21 These four sites rest on the same calcareous hard limestone bedrock, are positioned on southfacing slopes, experience similar temperatures (annual mean 17.7−19.1 °C) and are fenced to exclude large grazers. Experiments that were designed to simulate decreased (“drought”) and increased (“irrigation”) rainfall conditions took place at sites 3 and 4, starting 6 years before our sampling. At these two sites, five plots (25 m × 10 m each) were assigned for controls and five for each treatment. In the drought treatment plots, rain shelters, consisting of horizontal transparent plastic strips, decreased the amount of rainfall reaching the ground by ∼30% without affecting the rain frequency. In the irrigation treatment plots, water distributed through sprinklers, concurrently with rainfall, increased the water input to the soil by ∼30%. The water used in these manipulations came from the water grid, which is fed by local groundwater and by Lake Kinneret water (with soluble PO4 concentration on the order of 3 μg·L−1),22 with typical water δ18O values of −5‰ and 0‰, respectively. The total area of each site is ∼0.5 ha. To these established sites, we added an additional site (site 5) at the Golan Heights (32°58′0.66″ N, 35° 44′58.36″ E), which is also positioned on a south-facing slope and has a similar climate to that of site 4 but is located on basalt bedrock (Table 1). The soil type in this site is Nazzazic Vertisolic brown Mediterranean soil, which correlates with Vertic Luvisols in the FAO classification system and with Vertic Haploxeralfs in the USDA classification system. Soil was sampled from sites 2 and 3 four times during 2008: in January, April, August, and October, to cover the seasonal cycle. Soil samples from sites 1, 4, and 5 were taken once, during August 2010. At all sites, two samples per plot were taken from the entire soil profile (10−20 cm deep) and pooled together. Soil samples were always taken at the downslope side 2157

dx.doi.org/10.1021/es203551s | Environ. Sci. Technol. 2012, 46, 2156−2162


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of a shrub of the dominant vegetation type at the site. The dominant vegetation was Sarcopoterium spinosum (thorny burnet) in sites 1−4 and Alhagi graecorum Boiss at site 5. This was done in order to sample similar conditions and to focus on spots in which high biological activity takes place. Soil samples were dried at 40 °C, ground and sieved to 2 mm, and stored at 4 °C until analysis. This combination of dry soil and low temperature prevented resetting of the phosphate isotopic compositions, as was proved by repeated analysis, 6 months apart, of the same samples. 2.2. Analytical Methods. The δ18OP values of the resin-P were measured by a method recently developed at our lab.3 It is based on scaling up established methods for soil plant-available phosphate concentrations measurements,23 combined with methods for phosphate oxygen isotope analysis. Briefly, phosphate was first extracted from the soil by shaking it in deionized water with anion-exchange membranes (BDH-55164, from VWR Ltd.). The phosphate was eluted from the washed membranes by shaking in 0.2 M HNO3 overnight. The phosphate concentration in our dry (and rewet) soils was not higher than in the wet soils, indicating that, in contrast to other soils,24 in these soils there is no significant release of phosphate from microbial cell lysis during rewetting. Soil organic matter was removed from the extract by shaking it overnight with Superlite DAX-8 resin (Supelco/Sigma Aldrich). The phosphate was then extracted and purified by the precipitation of cerium-phosphate as described by McLaughlin et al.25 The precipitate was dissolved in nitric acid, the cerium was removed by a cation-exchange resin, and finally the phosphate was precipitated as silver phosphate. The δ18OP of the HCl-P was assessed following a recently developed method.2 This method is based on extraction with 1 M HCl and purification by Superlite DAX-8 resin and through successive precipitations of ammonium phosphomolybdate and magnesium ammonium phosphate. While this extraction also includes the resin-P fraction, the concentration of P in the HCl fraction is more than an order of magnitude larger, so the contribution of the available-P to this isotopic values is small. We have verified that hydrolysis of organic-P by HCl was negligible, by extraction with 18O-labeled HCl solution. Incorporation of labeled oxygen from the HCl solution to the PO4, which occurs during hydrolysis,26 was not observed. The phosphate concentration of all extractions was determined colorimetrically following the method of Murphy and Riley.27 For determination of isotopic composition, replicates of ∼0.3 mg of silver phosphate were packed in silver capsules and introduced into a high-temperature pyrolysis unit (HT-EA), where they were converted to CO in the presence of glassy carbon.28,29 The HT-EA is interfaced in continuous flow mode, through a GC column, to an isotope ratio mass spectrometer (Sercon 20−20). All isotopic values are given in the delta notation versus Vienna standard mean ocean water (VSMOW). The average standard deviation between replicates was 0.3‰. All measurements were performed against Ag3PO4 lab standards, which were calibrated against the TU-1 and TU-2 standards29 and against the IAEA-601 benzoic acid standard.30

Figure 1. Seasonal and between-treatments variability in resin-P δ18OP at (a) site 3 and (b) site 2.

δ18Op values is observed in January and the minimum variability in April. The highest (19.1‰) and lowest (14.5‰) δ18OP values at these two sites also occur in January. At site 2, in January, there are statistically significant lower values in irrigation plots in comparison to the values in drought plots (Figure 1b). The seasonal variability in resin-P concentrations at site 2 (Supporting Information Table S1) showed similar seasonality to that at site 3, already reported elsewhere,1 with the highest concentrations in January [5.6 ± 2.6 μg of P·(g of soil)−1 on average] and the lowest concentrations in April [2.9 ± 1.0 μg of P·(g of soil)−1]. The concentrations and δ18OP values of resin-P across all five sites are presented in Figure 2. The concentrations varied

3. RESULTS The results of seasonal sampling at sites 2 and 3 and across plots with control, drought, and irrigation treatments are presented in Figure 1. The resin-extractable phosphate δ18OP values vary in the range 14.5−20‰ for both sites and all treatments. For both sites the highest variability in resin-P

Figure 2. Variability in resin-P concentrations and δ18OP at all five sites. The δ18OP values are uncorrelated with the concentrations over an order-of-magnitude range. 2158



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between 1 and 12 μg of P·(g of soil)−1 for sites 1−4, while for site 5 the concentration range was 26−52 μg of P·(g of soil)−1. The δ18OP values of resin-P at all five sites were in the range 14.5−21.2‰ (only one sample, from site 4, had a value above 20‰). The δ18OP values of the HCl-P are presented in Figure 3. The low content of HCl-extractable phosphate at site 4

to reduce the uncertainty posed by using the literature values. By adopting the rainfall values to represent soil−water, we implicitly assumed that the biological activity in drying soils is low and most of the activity is in wet soils. Use of these values in eq 1 results in typical δ18OP values in the range of 28−40‰ for equilibration with leaf water. Organic-P synthesized by plants using this PO4 is expected to be heavy also. In contrast, the typical δ18OP values for equilibrium with soil−water are 19.5‰ in January and 16.0‰ in April. These values are expected to represent the δ18OP of phosphate equilibrated inside soil microbial cells and in plants stems (in contrast to leaves), which both have the same water isotopic composition. It can possibly also take place in the soil outside cells if active enzymes are present in this soil. The cellular P in microbes and plants is partly in inorganic and partly in organic form. For example, Bunemann et al.38 found that 39% of the P in the fungi they studied, and 13% of the P in the bacteria, were in orthophosphate form. If this inorganic P is released directly to the soil Pi pool, it will contribute phosphate with the isotopic values estimated above. In contrast, the cell organic-P needs to be mineralized by extracellular reactions, with their associated kinetic effects, before entering the soil Pi pool. The kinetic effects of the extracellular enzyme hydrolysis reactions reported thus far, are given by eq 2. Assuming that fractionation measured in lab experiments represents the processes in the soil, using the values we have calculated above for assumed formation of organic-P in equilibrium with intracellular water and based on the fact that APase is the dominate phosphatase in these alkaline soils,11 yields δ18OP values for inorganic-P of 4‰, −6‰, and 5‰ for phosphate mineralized from soil−water equilibrated RNA, DNA, and phosphomonoesters, respectively. From hydrolysis of leaf-water equilibrated organic-P with assumed δ18OP values of 28−40‰, we expect to get δ18OP values in the range 4−21‰. The upper range represents the case of organic-P equilibrated with highly evaporated leaf water and is probably less relevant in the winter and spring months when most biological activity takes place. Overall, these calculations predict that the mineralized inorganic-P will have a wide range of δ18OP values, which are usually considerably lighter than the values in equilibrium with soil−water. Similar predictions were made for marine systems, in which δ18OP values lower than that of equilibrium were expected for hydrolysis-derived phosphate.39 In contrast to this prediction for mineralized Porg, the observed δ18OP values of resin-P were in the range 14.5−20‰, which is close to the values we calculated above for equilibration with soil−water (16−19.5‰). This finding is consistent across a wide range of inorganic-P concentrations and across the seasonal cycle, rainfall manipulation experiments, and natural rainfall gradient of 90−780 mm annually. It is also consistent over a wide range of resin-P concentrations (Figure 2). In addition, only the higher end of the observed resin-P δ18OP values overlaps with that of the sedimentary phosphate, while similar resin-P values were found for soils with widely different δ18OP values of the primary mineral phosphate: 21− 22‰ for sites 1−4 and 6‰ for site 5. This suggests that the close-to-equilibrium values are not an accidental result of a mixture between light phosphate from hydrolysis and heavy phosphate from sedimentary apatite. The HCl-P δ18OP values in sites 2 and 3 were also close to those expected from equilibrium with the soil−water. This probably indicates that most of the inorganic-P at these sites already went through biological recycling. Values of total and HCl-extracted

Figure 3. The δ18OP of HCl-P at four of the sites.

prevented establishing a δ18OP value by our standard protocol. The isotope values obtained were in the range 17.1−21.3‰ at site 1, 15.6−17.5‰ for sites 2 and 3, and 8.2−10.9‰ for site 5.

4. DISCUSSION In order to analyze the results we have obtained, we first have to estimate the δ18OP values expected from mineralization of organic-P and from equilibrium effects, and from dissolution of primary phosphate minerals in the soil. Estimating the last term is relatively straightforward, since no isotopic fractionation is involved in the dissolution of apatite,16 which is the main Pbearing primary mineral. The δ18OP range of sedimentary phosphorites (not exposed to high temperatures) in the study area is 21−22‰,31 after adjustment to the revised value of the NBS120b standard. Hence, this is the expected range for primary-P in sites 1−4. In contrast, site 5 is located on a basalt bedrock, with δ18OP of ∼6‰.32,33 To estimate the δ18OP values expected from equilibrium fractionations, we have to take into account the temperature and isotopic composition of the water in which the equilibration takes place. The typical soil temperature at the sites, as well as the daily average air temperature, are around 10 °C in January and 17 °C in April and October. The higher temperature of ∼30 °C in August is considered here as mostly irrelevant, since it occurs when the soil is dry and the biological activity is minimal. The accumulated precipitation δ18O varies between ∼−4‰ in January to ∼−6‰ in April in the vicinity of site 3,34 and similar weighted mean precipitation δ18O of ∼−5‰ was observed for the areas of all the other sites (with no geographical variations).35,36 Leaf water is considerably enriched by evaporation, and a previous study in Israel reported δ18O values in the range from +5‰ to above +20‰.37 We did not measure directly the leaf and soil values, since they are known to change with time and with soil depth, and and without knowledge about the timing in which the equilibration takes place in the leaves and in each soil depth, it will be impossible 2159



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conserved in these samples as a result of the low annual rainfall (90 mm) at this arid site. The seasonal variations we observed also support the notion that the resin-P δ18OP values results from competition between different processes. We observed the highest variability in resinP δ18OP in January 2008 at both sites in which seasonal sampling took place (sites 2 and 3). The values cover the range from that expected for equilibrium at that month’s temperature (19.5‰) to values that are a few permil lower (down to 14.5‰). In contrast, the April samples show the lowest variability in phosphate oxygen isotopic compositions. These results are consistent with our suggestion above, of processes that erase the mineralization signal. In the year of the study, the rainy season started later than usual, and soil moisture started to increase only a few weeks before the January sampling. A previous study at these sites21 reported high soil respiration rates, probably associated also with fast mineralization rate, following the first rains of the season. This would result in a large flux of hydrolyzed Pi. In contrast, the soil microbial and protozoa activity may have still been low and spatially heterogeneous at the time of sampling. This combination may have caused the erasing of the mineralization signal to be only partial in some of the soil samples. These samples partially retain the light isotopic values resulting from the hydrolysis process. The higher biological activity in April, resulting from wet soil and higher temperature, resulted in values close to equilibrium in all samples. The changes in resin-P concentration can be also explained by high mineralization rates and low uptake rates early in the rainy season and higher plant and microbial uptake in spring, when plant activity is maximal.

inorganic-P, close to equilibrium, were also reported for other soils.2,40,41 We suggest here four possible processes that can contribute to the heavier than expected resin-P δ18Op values we observed: (1) It is possible that extracellular pyrophosphatase activity is causing equilibration with the soil−water, as suggested for dissolved δ18Op in the ocean.39 The soils we study go through wet−dry cycles that can release enzymes from inside the cells. Indeed, free pyrophosphatase was observed in dried soils that were rewetted in the lab.42,43 (2) The fractionation values we used for organic P mineralization are based on lab studies, which may not be relevant for the processes in soils. (3) Fractionation by microbial and plant uptake may cause an increase in the soil available inorganic δ18Op and may result in values that are accidentally close to the ones expected from equilibrium with soil−water. (4) The microbiota might release large amount of Pi which was equilibrated inside the cells (Figure 4). This flux may

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Figure 4. Suggested scheme of the main processes that control the oxygen isotopic composition of soil inorganic phosphate. Organic-P fluxes are shown as solid lines and inorganic-P fluxes as dashed lines. The phosphate is equilibrated within plants and microbial cells. Microbial cells, and protozoa feeding on it, excrete inorganic P. Plant litter and dead microbial cells contribute to the soil organic P pool. This pool is hydrolyzed with associated kinetic isotope effects. Equilibration of the inorganic pool by extracellular pyrophosphatase is also possible, as well as fractionation in uptake by microbiota and plants, protozoa. To simplify the scheme, we have omitted organic-P flux from protozoa, inorganic-P flux from plants, and uptake fluxes.

ASSOCIATED CONTENT

S Supporting Information *

The full resin-P data set, including concentration and oxygen isotope composition, appears in Table S1. This material is available free of charge via the Internet at http://pubs.acs.org.

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AUTHOR INFORMATION

Corresponding Author

*E-mail: [email protected]; Telephone: 972-2-6584758; Fax: 972-2-5662581.

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result from bacterial cell lysis,24 from Pi excretion by bacteria,44 or from protozoa that feed on bacteria and excrete surplus inorganic-P.45,46 More research is needed to determine the relative contributions of the various processes above to the values close to equilibrium we report. We suggest that such research should be based primarily on controlled incubations of soils in the lab, because separating the different effects in the field is difficult. It should be noted that if process 4 in the list above is proved to be the main one, then it will be possible to utilize the δ18OP to trace P turnover through the microbiota. The lower HCl-P δ18OP values at site 5 indicate that, at this site, a considerable fraction of the inorganic-P still reflects the primary apatite minerals. Using a value of 18‰ to represent the biologically equilibrated phosphate at this site (Figure 2) and 6‰ for the primary mineral, we can estimate that ∼75% of the total inorganic P in this soil is derived from primary apatite. The high values in two samples from site 1 resemble values for sedimentary apatite. This primary apatite was probably

ACKNOWLEDGMENTS This research was supported by ISF Grant 870/08. We thank Y. Kolodny for his comments on an earlier version of the manuscript and T. Venneman for providing Ag3PO4 standards. Four of the study sites are part of the GLOWA Jordan River project (www.glowa-jordan-river.de) supported by the German Federal Ministry of Education and Research (BMBF) in collaboration with the Israeli Ministry of Science and Technology (MOST). We also thank three anonymous reviewers for their constructive comments and suggestions.

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