Published January 30, 2014
The 11th Dahlia Greidinger Memorial Symposium:
Advanced Methods for Investigating Nutrient Dynamics in Soils and Ecosystems
Oxygen Isotopes for Unraveling Phosphorus
Transformations in the Soil–Plant System: A Review
Federica Tamburini*
Verena Pfahler
Christian von Sperber
Emmanuel Frossard
Institute of Agricultural Sciences
ETH Zurich
Research Station Eschikon 33
8315 Lindau, Switzerland
Stefano M. Bernasconi
Geological Institute
ETH Zurich
Sonneggstrasse 5
8092 Zurich, Switzerland
Phosphorus is a major nutrient for all living organisms. In the terrestrial environment, P is considered a double-edged sword. In some areas, agricultural
production is strongly limited by the low soil P availability, while in others,
P inputs in excess of plant needs have resulted in pollution of water bodies.
A better understanding of soil–plant P cycling is needed to provide agricultural and environmental managers with better concepts for P use. Together
with the routine analysis of soil available P, the determination of P chemical
forms, and the use of P radioisotopes, researchers have recently started using
the ratio of stable oxygen isotopes in phosphate (d18O-P). The scientific community interested in using this isotopic tracer is slowly but steadily expanding
because d18O-P has proven to provide important information on biological
processes influencing the P cycle and it could be used to trace the origin and
fate of P in soil–plant systems. This review examines the published results and
compiles the available data relevant for soil–plant systems, pinpoints gaps in
analytical techniques and knowledge, and suggests key questions and topics
to be investigated.
Abbreviations: d18O-P, ratio of stable oxygen isotopes in phosphate; Pi, inorganic
orthophosphate; Po, phosphorus-containing organic compounds; PPase, pyrophosphatase;
TCA, trichloroacetic acid; VSMOW, Vienna Standard Mean Ocean Water.
P
hosphorus is the 10th most abundant element on Earth and an essential
nutrient for all organisms. Phosphorus is fundamental to many biological
processes because it is involved in energy transfer and is the constituent of
a number of organic molecules (Westheimer, 1987). When in excess in the environment, however, P can become a pollutant, causing eutrophication of water
bodies (Sutton et al., 2013) and eventually important shifts in ecosystems. For all
these reasons, P chemistry and biochemistry, its cycle in marine, aquatic, and terrestrial environments, and its transfers from sources to sinks have been extensively
studied (Frossard et al., 2011; Paytan and McLaughlin, 2011; Ruttenberg, 2003).
Along with advances in technology, new analytical tools have provided deeper insights into P forms, pool size, transfers and fluxes, and processes affecting P cycling
(Frossard et al., 2012).
Stable isotopes have been used to track elements (i.e., O, N, S, and C) during
transfers between pools and to understand the respective roles of abiotic and biotic
processes during these transfers. Phosphorus can be bound to different elements
(e.g., C and O in phosphonates, N in phosphazenes, H in phosphine, different
elements in phosphides), but in soil–plant systems the vast majority of P is bound
to O, forming phosphate and to a lesser extent phosphonate and poly/pyrophosphate. This review shows whether, and under which conditions, the ratio between
the heaviest and lightest O isotopes in phosphate (18O and 16O; d18O in phosSoil Sci. Soc. Am. J. 78:38–46
doi:10.2136/sssaj2013.05.0186dgs
Received 20 May 2013.
*Corresponding author ([email protected]).
© Soil Science Society of America, 5585 Guilford Rd., Madison WI 53711 USA
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Soil Science Society of America Journal
phate) can be used to study P cycling and transformations. The
isotopic composition of O is reported in the conventional delta
notation in parts per thousand (denoted as ‰), calculated as
 Rsample

=
d 18O 
−1  1000
 Rstandard 
[1]
where R is the 18O/16O ratio in the sample and the standard, respectively. The d values are expressed relative to the international
standard Vienna Standard Mean Ocean Water (VSMOW). The
enrichment or depletion of the heavy isotope 18O between two
substances A and B is expressed by the fractionation factor a:
a A −B =
RA
RB
[2]
where RA and RB are the 18O/16O ratios in Samples A and B,
respectively. Because a is close to 1, the fractionation is commonly expressed as e, which is defined as
=
e
( a −1 ) 1000
[3]
Earlier research using this approach dated back to the 1940s,
when Winter et al. (1940) used 18O-labeled water to study the
exchange of O between water and several oxyanions, among
them phosphate. The outcome of this study was that O exchange
between phosphate and water measured in solution was slow
and not easily detectable under the experimental conditions (at
100°C for a 24–41-h reaction). Studies in biochemistry used the
measure of excess 18O in phosphate coming from labeled water
to understand the mechanism and the reaction rates of different
enzymes and to investigate biological processes, e.g., adenosine
triphosphate and the energy cycle (Boyer, 1978; Cohn, 1958).
In fact, as was observed later, only biological activity (e.g., P cycling within living organisms and cells and/or the effect of enzymes hydrolyzing organic P compounds) promotes O exchange
between water and phosphate at temperatures that are typical
for the Earth surface (Blake et al., 2005; Kok and Varner, 1967;
Longinelli and Nuti, 1973). For this reason, d18O-P was suggested as a possible “ideal life detection method” to be used on other
planets or on extraterrestrial material (Kok and Varner, 1967).
The resistance to breakage of the P–O bond under abiotic
conditions made d18O-P the perfect candidate to be used as a
paleotemperature proxy, together or as a substitute for oxygen
isotopes in carbonate minerals. Longinelli and Nuti (1973)
published the first empirical equation of the temperature dependence of the fractionation between phosphate and water, which
was later confirmed by Kolodny et al. (1983) and Shemesh and
Kolodny (1988). The equation was successfully used in paleoenvironmental studies to provide paleotemperature estimates from
bones, fish remains, and sedimentary phosphorites (e.g., Fricke
et al., 1998; Blake et al., 2010). Studies of marine and terrestrial
environments started in the 1980s (for an extensive review in the
field of marine studies, see Paytan and McLaughlin, 2011). In
this review, we focus on the application of d18O-P in soil–plant
www.soils.org/publications/sssaj
systems, with emphasis on methodological problems and propositions for future development.
Theory
The theory behind the use of d18O-P in environmental
studies has been extensively discussed elsewhere (Blake et al.,
2005; Paytan and McLaughlin, 2011). Therefore, we provide
only a short summary of the main factors controlling the O isotope composition of phosphate in nature:
1. Oxygen exchange between phosphate and water is slow
and negligible in abiotic systems under Earth surface
conditions (Blake et al., 2005; Lécuyer et al., 1999; Winter
et al., 1940), thus O will preserve its original isotopic
composition in the absence of biological activity.
2. Processes such as adsorption–desorption and
precipitation do not produce any significant O isotope
fractionation (Jaisi et al., 2010; Liang and Blake, 2007). This
has been recently confirmed for sorption–desorption by the
use of the radioisotopes 32P and 33P (Randriamanantsoa et
al., 2013). To date, however, there are no studies showing
whether partial dissolution of minerals by organic acids in
soil–plant systems (e.g., oxalate or other organic compounds
present in soil and secreted by microbes and/or plant roots)
would cause any fractionation.
3. It has been reported that microorganisms preferentially
take up lighter isotopologues of phosphate, leading to
an enrichment of heavier isotopologues in the residual
phosphate. However, this effect was only observed under
laboratory conditions and using one organism, E. coli,
which showed a fractionation factor e of −3‰ (Blake et
al., 2005). So far, no information is available on the effect
of uptake by other bacteria, plants, or other organisms.
4. Inside living cells, P is involved in several biochemical
processes (Westheimer, 1987). The ubiquitous
intracellular enzyme pyrophosphatase (PPase) catalyzes
the hydrolysis of pyrophosphate (P2O7). During the
hydrolysis of P2O7, one O atom from the surrounding
water is incorporated into the P2O7 molecule, leading
to the formation of two inorganic orthophosphate
molecules, which are subsequently released. This process is
extremely important for the living cell because it controls
the concentration of P2O7, which otherwise would rise to
toxic levels (Cooperman et al., 1992). Pyrophosphatase
enzymes lead to a complete O exchange between water
and phosphate, even in the absence of P2O7 (Cohn, 1958;
Blake et al., 2005). This observation has been attributed to
the fact that not only P2O7 molecules but also phosphate
can bind at the active site of PPases. After the binding
of two phosphate molecules, enzyme-bound P2O7 is
formed, which immediately is hydrolyzed again, releasing
phosphate. The formation of enzyme-bound P2O7 and its
subsequent hydrolysis, which includes the incorporation
of one O atom from water, is extremely fast. Therefore, this
process leads to a complete O exchange between water and
phosphate with time. This enzyme-catalyzed O exchange
is subject to a thermodynamic isotopic fractionation
39
leading to a temperature-dependent equilibrium between
water and phosphate, which has been described by the
empirical equation of Longinelli and Nuti (1973):
T=
111.4 − 4.3 ( d 18O-P − d 18O-H 2O )
[4]
where T is the temperature (°C) and d18O-P and d18OH2O are the O isotope composition of phosphate and water (‰ with respect to VSMOW). Other enzymes, such
as ATPase, also promote the exchange of all O atoms in
phosphate (Webb and Trentham, 1981). However, the
isotopic fractionation associated with these enzymes has
not yet been determined. The ubiquity of PPase in living
cells probably overwrites any other fractionation effects
on the d18O of intracellular phosphate. Therefore, once a
phosphate molecule has entered a living cell, its O isotopic
composition will be in isotopic equilibrium with cell water (Paytan et al., 2002, Blake et al., 2005).
5. The hydrolysis of phosphoesters by enzymes
(phosphatases or other enzymes that release phosphate
as byproduct, such as nucleotidases; Liang and Blake,
2006b) will release phosphate with a different isotopic
composition than the original organic compound.
Depending on whether the enzyme catalyzing the
reaction is a phosphomonoesterase or phosphodiesterase,
hydrolysis of phosphoesters leads to the cleavage of one or
two P–O bonds. In the case of phosphomonoesterases,
the released phosphate inherits three O atoms from the
original organic compound, while one is incorporated
from a water molecule. In the case of phosphodiesterases,
the released phosphate inherits two O atoms from the
original organic compound and two O atoms from water
molecules, one by the phosphodiesterase and the other by
the phosphomonoesterase hydrolyzing the new-formed
monoester. The incorporation of O from water is subject
to an isotopic fractionation, which depends on the type
of enzyme and substrate. At present, only a few enzymes
have been studied: alkaline and acid phosphatase,
5¢-nucleotidase, DNase, and RNase, with fractionation
factors (e) ranging between −30 and 20‰ (Liang
and Blake, 2006b, 2009; von Sperber et al., 2013). No
information is currently available on other enzymes (e.g.,
phytase) that are highly relevant for the soil–plant system
(Nannipieri et al., 2011).
Analytical and Preparation Methods
One of the most challenging aspects of this type of research
has been the preparation of samples and their analysis. The first
studies used excess 18O, which was determined by mass spectrometry (Cohn, 1958), by optical emission of CO+ ( Johansen et al.,
1990; Larsen et al., 1989), or by 31P nuclear magnetic resonance
(Cohn and Hu, 1978). Only later was the d18O in phosphate at
natural abundance reported. The first techniques developed for
isotopic analysis involved the isolation and purification of PO4
as BiPO4 and used a fluorination method (i.e., reaction of BiPO4
with BrF5) to release O2 for analysis (Lécuyer, 2004). Bismuth
40
phosphate is a relatively hygroscopic compound (Tudge, 1960),
which makes the further analysis of O isotopes difficult due to
possible contamination by absorbed water. The use of Ag3PO4
was first introduced by Firsching (1961), and then proposed
again by Crowson and Showers (1991). The choice of Ag3PO4
made the isolation of P from natural apatites and the purification
procedure simpler than the protocol proposed by Tudge (1960),
and, most importantly, Ag3PO4 was not hygroscopic; however,
the purification procedure used to obtain Ag3PO4 needs to be
adapted to the different nature and chemistry of the samples to
be analyzed. This means that there cannot be a single unified
approach to isolate and purify P from a sample. Because of the
extensive manipulation required by most of the published purification procedures, the Ag3PO4 produced is always <100%
of the phosphate present in the initial extract; however, no fractionation due to incomplete recovery of the phosphate has been
observed (Tamburini et al., 2010; Colman, 2002).
Several methods have been used to analyze O isotopes in
phosphate following technical advancements in mass spectrometry. Vennemann et al. (2002) provided an extensive comparison
and discussion of these methods. Recent improvements in isotope
ratio mass spectrometry (Brand, 1996) and the use of high-temperature reduction (temperature conversion elemental analysis, TCEA, where Ag3PO4 is converted quantitatively to CO at 1450°C)
coupled to a continuous-flow mass spectrometer has led to important improvements, specifically a reduction in both the sample size
and the time required for analysis. The main issue of this analytical
method, however, is the need for routinely measuring replicates of
standards and samples in each run because changing pyrolysis conditions can strongly affect the measurements (Tamburini et al., 2012a;
Vennemann et al., 2002). Unfortunately no internationally certified
standards for Ag3PO4 are currently available, despite Ag3PO4 being
considered the most suitable standard for measurements of 18O in
both phosphate and other O-containing compounds (Fourel et al.,
2011). Finally, Melby et al. (2011) proposed the use of electrospray
ionization mass spectrometry (ESI-MS) for measuring the relative
isotopic composition of inorganic P enriched in 18O (i.e., the relative abundance of P18O16O3, P18O216O2, P18O316O, and P18O4)
in liquid samples.
Extraction and Purification Methods for Soils
Soil P is distributed in many inorganic and organic pools
(Frossard et al., 2011), which are interconnected by a series of biotic and abiotic reactions (Fig. 1). Soil P pools have been largely
studied by the Hedley sequential fractionation scheme (Hedley
et al., 1982), which extracts soil inorganic and/or organic P (Pi
and Po, respectively) successively with an anionic resin (plantavailable Pi), NaHCO3 (Pi and Po loosely sorbed on clays and
mineral particles), NaOH (Pi and Po associated with Fe and Al
oxides and Po), and HCl (mineral P associated with Ca). The
phosphate pools obtained by the Hedley fractionation are operationally defined, but they give a fair idea of the P forms present in
a soil (Negassa and Lainweber, 2009). At present, the analysis of
the d18O in soil phosphate pools has targeted the resin-extractSoil Science Society of America Journal
Plant residues
Organic fertilizers
P exports
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SLOWLY EXCHANGEABLE
Pi
RAPIDLY EXCHANGEABLE
Pi
The presence of these extra organic
compounds
complicates the O isotope
MIneral fertilizers
analysis of phosphate. Because the isotopic signature of Po will partially deterRoot
uptake
mine the signature of the released phosDesorption/
MICROBIAL
phate (Liang and Blake, 2006b), however,
solubilization
Immobilization
PHOSPHATE IN THE
BIOMASS
measuring the d18O-P of organic comSOIL SOLUTION
Micro and
Adsorption/
Mineralization
macro fauna
pounds would be of great importance. In
precipitation
P-limited environments, where phosphate
concentrations are low, Po compounds
Death
Assimilation
Enzymatic
Biological
are considered to be an important source
hydrolysis
solubilization
of phosphate for soil organisms and plants
RAPIDLY MINERALIZABLE
Po
(Nannipieri et al., 2011).
Despite this, there is still a lack of a
SLOWLY AND NOT
suitable procedure to purify phosphate
MINERALIZABLE Po
from extracts targeting the Po pool.
Erosion/runoff
Organic compounds are generally quanLeaching
Weathering of
tified after total digestion of the extracts
parent material
at temperatures >100°C (Hedley et al.,
Figure
1
Fig. 1. Schematic representation of the P cycle in soil–plant systems, including organic and 1982). This approach, however, is inapinorganic P (Po and Pi, respectively) (modified after Frossard et al., 2011).
plicable for the extraction of Pi from Po
and the measurement of d18O-P because
able Pi (Weiner et al., 2011), the HCl-extractable P (Tamburini
a reaction temperature >70°C together with the use of strong acet al., 2010), and the microbial P (Tamburini et al., 2012b).
ids will cause the exchange of O between phosphate and the soluZohar et al. (2010a) combined the Hedley sequential extraction
tion. Liang and Blake (2006a) reviewed several methods for anawith the purification protocol of McLaughlin et al. (2006c) to
lyzing the d18O-P of organic compounds. They concluded that
18
get the d O-P of inorganic phosphate from different soil pools.
digestion with ultraviolet (UV) radiation is the best approach,
Some of the procedures for purifying inorganic phosphate
causing little fractionation (Liang and Blake, 2006a); however,
extracted from soils (Markel et al., 1994; Tamburini et al., 2010)
this has not yet been applied to a soil system.
and also from sediments ( Jaisi and Blake, 2010) follow the puriExtraction and Purification Methods for Plants
fication scheme used for apatites, described by Tudge (1960) and
Kolodny et al. (1983). A series of precipitations, first ammonium
Phosphorus is also present in different forms in plants: free
phosphomolybdate (at low pH) and then magnesium ammoorthophosphate, nucleic acids, RNA, DNA, phospholipids, and
nium phosphate (at high pH) purifies the extracts and separates
other ester P (Veneklaas et al., 2012). The O isotope signature of
phosphate from other inorganic and organic compounds. Other
these different P forms would be of great interest for soil–plant
procedures (McLaughlin et al., 2006c; Weiner et al., 2011;
system studies because plant material constitutes an important P
Zohar et al., 2010a) use cerium phosphate to separate and purify
source in soils. To date, the analysis of d18O in phosphate from
phosphate from the extracts. This approach was first designed to
plant material has been reported in only three studies. Young et
process marine water samples, and cerium phosphate was chosen
al. (2009) used water to extract phosphate from vegetation sambecause cerium does not form any salt with Cl (McLaughlin et
ples and purified it with the McLaughlin protocol. Tamburini et
al., 2004, 2006c). This purification method, however, is not alal. (2012b) analyzed a range of plant species sampled in a glacier
ways applicable to soil extracts that have high concentrations of
forefield using a diluted trichloroacetic acid (TCA) extraction.
organic compounds and oxides (Tamburini et al., 2010). The puPfahler et al. (2013) used a two-step extraction targeting inorrification steps are of great importance because several ions (e.g.,
ganic and organic phosphate from soybean [Glycine max(L.)
Na and Cl) hinder the precipitation of Ag3PO4, and the presMerr.] leaves and seeds. A diluted TCA extraction was chosen
ence of other chemical species containing O (such as nitrates,
in these latter two studies because TCA extracts orthophosphate
sulfates, and organic compounds) in the final sample interferes
(Hawkins and Polglase, 2000) and only a few other organic
18
with analysis of the d O of phosphate. Tamburini et al. (2010)
compounds, such as sugar phosphates and phytate. Sugar phosshowed that extraneous O-containing compounds strongly afphates and phytate are not hydrolyzed during the extraction, so
fect the isotopic measurement of Ag3PO4. They suggested that
the d18O of phosphate from a TCA extract predominantly repthe pH adjustments of the sample extracts needed for the preresents the signature of phosphate. Trichloroacetic acid extractcipitation of cerium phosphate (optimum pH is 5–6) are responable P is sometimes referred to as metabolic P. Concentrated
sible for the precipitation of extraneous oxides and further adHNO3 was subsequently used to extract and hydrolyze the resorption of organic compounds on the oxides.
maining P pools, which mainly consisted of phospholipids and
41
other organic compounds. This extraction followed the protocol
described by Tudge (1960) and modified by Liang and Blake
(2006a). Both extractions were systematically performed using
18O-labeled and unlabeled acids to track the exchange of O between phosphate and the solution due to hydrolysis. The extracted solutions were then purified using the protocol described by
Tamburini et al. (2010).
Investigation of Phosphorus
Dynamics in Soil–Plant Systems
Because there are only a limited number of soil–plant system
studies using d18O-P, we have also referred to studies outside of the
field of the soil–plant system when their evidence might prove useful.
The first studies using O isotopes to study phosphate in
soil–plant systems were performed at the end of the 1980s
by Saaby-Johansen, Larsen, and Middelboe, who performed a
series of experiments using soils, plants, and soil microorganisms (Larsen et al., 1989; Johansen et al., 1990; Middelboe
and Saaby, 1998). They used KH2PO4 that was labeled with
both 18O and 32P to test if 18O could be used as a substitute
for radiolabeled phosphate in environmental studies. During a
period of 6 mo, at ambient temperature, they could not observe
a significant 18O delabeling of the added KH2PO4 in sterile
soils across a wide range of pH; however, they did observe that
the 18O label was lost after 5 wk when plants were grown on
these soils (Larsen et al., 1989). They estimated that 5% of excess 18O was lost for every millimole of CO2 emitted from biologically active soils ( Johansen et al., 1990). Their conclusion
was that 18O in phosphate could not be used as a substitute
for P radioisotopes but that it could give additional information on biological activity. They tested this hypothesis in a later
study, where they also proposed an equation relating bioactivity in soils with the loss of label during the experiment time
(Middelboe and Saaby, 1998).
A few years after the first experiments of Larsen and his
group, Ayliffe et al. (1992) and Mizota et al. (1992) used 18O-P
at natural abundances and analyzed environmental samples from
island phosphate deposits from the Pacific Ocean and volcanic
ash soils in Indonesia, respectively. The aim was to study the process of soil phosphate formation (the reaction between P-rich
guano and carbonates) and its origin (apatite from rocks vs.
biogenic material, such as bones). These studies showed that the
d18O-P was probably influenced by biological activity. Ayliffe et
al. (1992) postulated that the action of soil microorganisms, by
cycling phosphate from guano, would reset the d18O-P to equilibrium, which is controlled by the d18O of water and the soil
temperature (see Eq. [4]).
One of the original aims of using d18O-P in soil–plant
studies was to identify sources of P and to trace its movement in
the environment (Melby et al., 2013). This idea was pursued in
several studies, which, however, showed results that were not encouraging. Indeed, the observed O isotope effects were strongly
dependent on local conditions. The first attempt was made by
Markel et al. (1994), who combined grain size characterization
42
of the sediments in Lake Kinneret, P pool characterization and
quantification (adsorbed P, P associated with Al, Fe, and calcite,
and P found in apatite), and the d18O-P in the grain size fractions
(clay, silt, and sand). From this, they could conclude that apatite
of a detrital origin was the main source of P to the lake, although
it did not play a significant role in biological productivity.
The study of Gruau et al. (2005) was also not particularly
encouraging with respect to the use of d18O-P as an environmental tracer. They characterized several fertilizers and compared the
results to the isotope signature of phosphate in recycled wastewaters. They observed that the d18O-P variability in the fertilizers was linked to the variability of the primary sources (e.g., the
phosphorite deposits from which the fertilizers were manufactured). They concluded, however, that the d18O-P of the sources
was not distinct enough to allow its use as a tracer. Furthermore,
they hypothesized that the d18O-P was probably influenced by
biological activity in the water, making it difficult to conclude if
the d18O-P in water was the primary isotopic signal of the source
or if it was the product of intense biological P cycling.
Regardless of these results, Paytan and co-authors used
d18O-P to trace P in lacustrine (Elsbury et al., 2009), estuarine
(McLaughlin et al., 2006a), and coastal systems (McLaughlin
et al., 2006b). They were able to distinguish between waters of
different origin and to recognize the effect of diagenetic processes. In this context, Young et al. (2009) published the first
collection of d18O-P values measured in rocks, phosphorites,
fertilizers, and waters.
Dust-borne P (phosphate sorbed on particles and brought
by aeolian material) can constitute an important source of P to
terrestrial and aquatic ecosystems (Okin et al., 2004), and d18OP has been used to trace atmospheric P inputs to Lake Kinneret
in Israel (Gross et al., 2013). Gross et al. (2013) sampled several
P sources around Lake Kinneret and compared their resin-extracted d18O-P with that extracted from dust samples collected
close to the lakeshores. They were successful in discriminating
between local P sources (from soils surrounding the lake) and
P coming from remote desert areas and brought to the studied
region by exceptional wind events.
Because of the strong impact of biotic processes on the
O isotope composition of phosphate, the most interesting
advancements have come from studies using d18O-P as an indicator of biological activity in soils and the environment. As
mentioned above, Larsen et al. (1989) were the first to realize
that biological activity was responsible for the d18O-P changes
observed with time in soil–plant systems. Their findings were
recently confirmed by Melby et al. (2013), who observed that
the d18O of P extracted by a modified Bray solution had disappeared 50 d after the addition of 18O-enriched phosphate
in incubated nonsterile soils, whereas it remained constant in
sterile soils. They calculated half-life values of 15 to 22 d for
the decrease of 18O in phosphate. Zohar et al. (2010b) were
the first to combine a sequential extraction of P with O isotope
measurements in the extracted phosphate from a series of soil
incubations treated with reclaimed wastewater and fertilizers.
Soil Science Society of America Journal
They also concluded that biological activity was changing the
isotopic signature of phosphate in the soil P pools, even the
ones generally considered to be more stable (e.g., P bound to
mineral phases extracted by NaOH or HCl).
In two studies, Angert et al. (2011; 2012) concentrated
their efforts in understanding d18O-P variability linked to rainfall and bedrock gradients. They looked at resin-Pi, as an indicator of readily bioavailable P, and at HCl-extractable Pi. As expected, this last pool, representing P bound to apatite, showed
values close to the parent material (Angert et al., 2012) and no
variability through the seasons (Angert et al., 2011). In contrast,
the d18O of resin-P changed with time, with values close to the
equilibrium as predicted by the Longinelli and Nuti equation
(Eq. [4]). They suggested that the rate of change was dependent
on the rate of biological activity in the soil, which was controlled
by season and humidity.
Another study by Tamburini et al. (2012b) measured d18O
in four P pools (microbial, resin- and HCl-extractable soil P
pools, and TCA-extractable P from plant material) from a soil
chronosequence in a glacial forefield to understand P cycling
during the early stages of soil development on a granitic parent material. They showed that mineral P (HCl-extractable Pi)
retained the O isotope signature of the original bedrock for
150 yr, while in older soils the d18O increased to reach values
similar to that observed in the resin-extractable and microbial
Pi. Furthermore, they showed that the TCA-extractable P in
plants was more enriched in 18O than the bedrock P source (as
also shown by Pfahler et al., 2013). In contrast, the d18O of the
microbial and resin-extractable Pi pools was close to the values
expected at equilibrium with soil water, underscoring the dominant effect of microbial activity even at the youngest soil site.
From these studies it is clear that soil biological activity leading
to equilibration with water is a dominant factor determining the
d18O of phosphate in soil pools.
Apart from Tamburini et al. (2012b), only two other studies have analyzed the d18O of phosphate in plants. Young et al.
(2009) investigated the isotopic signature of vegetation to determine the relevant isotopic differences among P sources to the
soil. They found differences among the investigated species and
different growth stages, although they did not attempt to identify the processes responsible for the d18O-P in plants. Pfahler
et al. (2013) were the first to study the effect of plant growth on
d18O-P. They used hydroponic cultures of soybean and analyzed
the d18O of TCA-extractable P, the d18O of leaf water, the d18OP of the source phosphate, and phosphatase activity in leaves. As
hypothesized, the d18O-P in the leaves showed a distinct signature different from the source phosphate, with values close to the
equilibrium with leaf water. Importantly, the rather high d18OP values observed in the TCA extracts of leaves were similar to
those observed by Tamburini et al. (2012b) in other plant species. Finally, besides measuring the d18O-P of TCA-extractable
P, Pfahler et al. (2013) also analyzed the d18O-P of the structural
P of the leaves (defined as the P that remains in the leaf material
once the TCA-extractable P has been removed). They found a
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very high 18O enrichment, reaching up to 60‰; however, they
were unable to explain this finding.
Considerations on the Natural Variability of the
Ratio of Stable Oxygen Isotopes in Phosphate
In Table 1 and Fig. 2, we present all the available d18OP data relevant for the soil–plant system. Phosphate extracted
from different sources (phosphorites, phosphate in plants, soil
P pools) shows d18O-P values spread across a relatively narrow
range, from 10 to 25‰. Only phosphate from igneous and volcanic rocks (High T rocks in Fig. 2) and structural P present
in plants plotted significantly outside this range. A possible
reason for the observed spread of data is that once taken up
by organisms—either plants, microorganisms, or animals—the
O in phosphate is partially or totally exchanged with the O in
water. Because of the narrow range of both O isotope values
in water and of the temperature under ambient conditions,
the O in phosphate is brought to a predictably narrow range
of temperature-dependent equilibrium values. In fact, only
phosphate from either igneous rocks (where apatite crystallizes
at high temperature) and from fertilizers derived from those
rocks (Gruau et al., 2005) are different from this range. The
reason why structural P in plants shows such heavy d18O is still
unclear (Pfahler et al., 2013).
Future Developments
Here we highlight areas of research that need further development and investigation, and suggest avenues to be explored
where the d18O-P tool could yield new and relevant information
on P transformations in soil–plant systems.
1. Process level: To achieve a better understanding of P
cycling, it is necessary to assess the fractionation induced
by more enzymes and most importantly by phytase because
this enzyme is very important in soil–plant systems. Along
with phytase, it is also important to assess the fractionation
induced by enzymes such as ligases that bind P to organic
compounds in order to understand, among other things,
why plant structural P is enriched in 18O.
2. Soil level: It has been shown by several studies (Angert
et al., 2012; Larsen et al., 1989; Tamburini et al., 2012b;
Zohar et al., 2010b) that at some point in time the d18O-P
of soil P compartments will reach equilibrium with water.
A better understanding of the kinetics of equilibration
between these pools may provide important information
on P cycling in soils. Using the approach developed by
Larsen et al. (1989), we could add highly 18O-enriched
P compounds (either artificially produced or naturally
enriched compounds such as structural P leaf litter) and
follow the rate of delabeling to obtain information on soil
biological activity and on organic P turnover, for which
there is still little information available. To achieve this,
it is also necessary to develop a method for purifying soil
extracts to correctly analyze the d18O-P of Po compounds.
Only then will we have a proper inventory of the d18O-P
of the different Pi and Po pools in the soil (even though
43
Table 1. Summary of the available 18O-P data on P sources (lithogenic and organic material and industrial products), soil P
pools, and phosphate in waters. Isotope data are ordered by increasing minimum values.
Type
Material
carbonatite
hydrothermal
volcanic ash
silt
sand
clay
aerosol (Israel)
phosphate rock
phosphorite (Florida)
phosphorite (Morocco)
dust (Israel)
calcarenite
Santa Barbara sediment
HCl-P
Soil pools
HCl-P
NaOH-P
HCl-P
Acetic acid-P
Bray-P
H2O-extractable P
Microbial P
HNO3–P
Resin-P
NH4F-P
Resin-P
Resin-P
HCl-P
Resin-P
HNO3–P
NaHCO3–P
Phosphate in organic material vegetation (TCA‡ extract)
vegetation (H2O extract)
animal feces
soybean leaves (TCA extract)
ivory and skulls
guano
humus
plankton
algae
soybean leaves (HNO3 extract)
detergent
Industrial products
fertilizers
fertilizers
toothpaste
fertilizers
marine waters (San Francisco Bay)
Waters
wastewater treatment plant
river and groundwater
Lake Erie waters (surface)
river water (California)
discharge water (France)
open ocean (Pacific)
† With respect to Vienna Standard Mean Ocean Water.
‡ TCA, trichloroacetic acid.
Lithogenic material
this will remain a broad characterization).
3. Plant level: Studies targeting the effects of different
parameters such as atmospheric humidity and the supply
of P, water, and CO2 should allow us to understand more
precisely how they affect the d18O-P of metabolic P and
structural P. It will also be necessary to show whether the
equilibrium equation of Longinelli and Nuti (Eq. [4]),
which was developed using phosphate from fish remains,
still holds true for plants and soil organisms.
44
Min. 18O-P
Max. 18O-P
———— ‰† ————
0.2
10
2.4
12.2
5.3
6.2
8.6
16.9
9.1
12.8
12.7
22.5
14.2
24.9
16.7
19.1
17.2
23.2
18.5
20.5
19.5
22.6
20.1
22.8
21
5.6
17.8
6
15.1
7.8
22.5
8.2
21.3
10.1
20.6
10.6
24.3
10.7
24.5
11.2
17.8
11.7
12.7
19.8
12.7
24.8
14.5
19.1
14.5
20
16.1
16.9
17.4
22.8
19
27
19
23.7
12.4
31.4
14.2
23.1
15.7
18.3
16.9
27.5
18.6
24
19.8
23.1
21.8
22.9
23.4
27.6
28.6
42.6
57.1
13.3
18.6
14.8
27
15.5
22.3
17.7
19.6
23.1
7.8
20.1
8.4
14.2
9.2
16.4
9.7
17.1
14.1
20.3
16.6
18.4
18.6
22
Reference
Mizota et al., 1992
Mizota et al., 1992
Mizota et al., 1992
Markel et al., 1994
Markel et al., 1994
Markel et al., 1994
Young et al., 2009
Ayliffe et al., 1992
Gruau et al., 2005
Gruau et al., 2005
Gross et al., 2013
Ayliffe et al., 1992
McLaughlin et al., 2006c
Zohar et al., 2010b
Tamburini et al., 2012b
Zohar et al., 2010b
Angert et al., 2012
Mizota et al., 1992
McLaughlin et al., 2006a
Zohar et al., 2010b
Tamburini et al., 2012b
McLaughlin et al., 2006c
Tamburini et al., 2012b
Mizota et al., 1992
Angert et al., 2011
Angert et al., 2012
Angert et al., 2011
Gross et al., 2013
McLaughlin et al., 2006a
Zohar et al., 2010b
Tamburini et al., 2012b
Young et al., 2009
Young et al., 2009
Pfahler et al., 2013
Mizota et al., 1992
Ayliffe et al., 1992
Ayliffe et al., 1992
McLaughlin et al., 2006c
McLaughlin et al., 2006c
Pfahler et al., 2013
Young et al., 2009
McLaughlin et al., 2006a
Young et al., 2009
Young et al., 2009
Gruau et al., 2005
McLaughlin et al., 2004, 2006b
Young et al., 2009
Young et al., 2009
Elsbury et al., 2009
McLaughlin et al., 2006a
Gruau et al., 2005
McLaughlin et al., 2004
4. Ecosystem level: Taking into account all the discussed
studies, it would appear that if applied in P-limited systems,
the d18O-P technique could provide information on
biological processes. When P is abundant and not limiting,
however, the d18O-P technique could be used to identify P
sources and sinks and trace P transfers through the system.
Provided the availability of P sources having distinct 18O-P
signatures, it would be of interest to trace their fate in the enviSoil Science Society of America Journal
Fig. 2. Ranges of d18O-P values of data presented in Table 1. Soil P
pools, as operationally defined, are grouped to give an approximate
idea of their form in soils: soil solution + easily exchangeable
P(H2O-P, resin-P, and NaHCO3–P); slowly exchangeable inorganic P
(Pi) (HCl-P, acetic acid-P, Bray-P, NH4F-P, and NaOH-P); and total P
(HNO3–P). Animal feces also include guano. Vegetation (structural)
represents HNO3–extractable P, while metabolic represents H2O- and
trichloroacetic acid-extractable P. Detergents include also toothpaste,
and High T rocks represent d18O-P from lithogenic material formed at
high temperature (e.g., igneous and volcanic rocks).
ronment so as to identify the P origin and/or a source of pollution. In this respect, the database started by Young et al. (2009)
and continued here (Table 1; Fig. 2) should be further developed, e.g., by exchanging data obtained in different laboratories.
To make the intercomparability of results of such a data set more
meaningful, it would be of great benefit for the entire community to develop internationally certified standards for isotope
measurements (Tamburini et al., 2012a).
Acknowledgments
We would like to thank A. Oberson for her help and fruitful discussion,
S. Granger, and two anonymous reviewers for their comments,
which improved the original manuscript; F. Tamburini is grateful
to the Greidinger family and to the BARD fund for sponsoring her
participation to the Dahlia Greidinger Symposium in 2013.
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