Advances in agronomy volume 125

ADVANCES IN AGRONOMY
Advisory Board

PAUL M. BERTSCH

RONALD L. PHILLIPS

University of Kentucky

University of Minnesota

KATE M. SCOW

LARRY P. WILDING

University of California, Davis

Texas A&M University

Emeritus Advisory Board Members

JOHN S. BOYER

KENNETH J. FREYw

University of Delaware

Iowa State University

EUGENE J. KAMPRATH

MARTIN ALEXANDER

North Carolina State
University

Cornell University

Prepared in cooperation with the
American Society of Agronomy, Crop Science Society of America, and Soil
Science Society of America Book and Multimedia Publishing Committee
DAVID D. BALTENSPERGER, CHAIR
LISA K. AL-AMOODI

CRAIG A. ROBERTS

WARREN A. DICK

MARY C. SAVIN

HARI B. KRISHNAN

APRIL L. ULERY

SALLY D. LOGSDON
w

deceased


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CONTRIBUTORS
Dominique Arrouays
INRA, InfoSol Unit, Orleans, France
Ruth E. Blake
Department of Geology and Geophysics, Yale University, New Haven, Connecticut, USA
Guanglong Feng
State Key Laboratory of Desert and Oasis Ecology, Xinjiang Institute of Ecology and
Geography, Chinese Academy of Sciences, Urumqi, China
Bruno Gerard
International Maize and Wheat Improvement Centre (CIMMYT), El Batan, Mexico
Michael G. Grundy

CSIRO, EcoSciences Precinct, Dutton Park, Queensland, Australia
Alfred E. Hartemink
University of Wisconsin-Madison, Department of Soil Science, Madison, USA
Ji-Zheng He
State Key Laboratory of Urban and Regional Ecology, Research Center for EcoEnvironmental Sciences, Chinese Academy of Sciences, Beijing, China, and Melbourne
School of Land and Environment, The University of Melbourne, Parkville, Victoria,
Australia
Jonathan W. Hempel
United States Department of Agriculture, Natural Resources Conservation Service, Lincoln,
Nebraska, USA
Gerard B.M. Heuvelink
ISRIC—World Soil Information, Wageningen, Netherlands
S.Young Hong
National Academy of Agricultural Science, Rural Development Administration, Suwon,
South Korea
Hang-Wei Hu
State Key Laboratory of Urban and Regional Ecology, Research Center for EcoEnvironmental Sciences, Chinese Academy of Sciences, Beijing, China, and Melbourne
School of Land and Environment, The University of Melbourne, Parkville, Victoria,
Australia
Deb P. Jaisi
Department of Plant and Soil Sciences, University of Delaware, Newark, Delaware, USA
Mangi L. Jat
International Maize and Wheat Improvement Centre (CIMMYT), NASC Complex, Pusa,
New Delhi, India

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x

Contributors

Xiangbin Kong
The College of Resources and Environmental Science, China Agricultural University, and
Key Laboratory of Farmland Quality, Monitoring and Control, National Ministry of Land
Resources, Beijing, China
Dinesh Kumar
Division of Agronomy, Indian Agricultural Research Institute, Pusa, New Delhi, India
Philippe Lagacherie
INRA, IRD, Lab Etud Interact Sols Agrosyst Hydrosyst, Montpellier, France
Rattan Lal
Carbon Management and Sequestration Center, The Ohio State University, Columbus,
Ohio, USA
Glenn Lelyk
Agriculture and Agri-Food Canada, University of Manitoba, Winnipeg, Manitoba, Canada
Baoguo Li
The College of Resources and Environmental Science, China Agricultural University, and
Key Laboratory of Farmland Quality, Monitoring and Control, National Ministry of Land
Resources, Beijing, China
Kejiang Li
Institute of Dryland Farming, Key Field Scientific Observation Station of Hengshui Fluvoaquic Soil Ecology Environment, Ministry of Agriculture, Hengshui, China
Hongbin Liu
Ministry of Agriculture Key Laboratory of Crop Nutrition and Fertilization/Institute of
Agricultural Resources and Regional Planning, Chinese Academy of Agricultural Sciences,
Beijing, China
Alexander B. McBratney
Faculty of Agriculture and Environment, The University of Sydney, Sydney, New South
Wales, Australia
Neil J. McKenzie
CSIRO Australia, Campus International de Baillarguet, Montpellier, Cedex, France

Maria d.L. Mendonca-Santos
EMBRAPA-Brazilian Agricultural Research Corporation/The National Centre of Soil
Research (Embrapa Solos), Rio de Janeiro, Brazil
Budiman Minasny
Faculty of Agriculture and Environment, The University of Sydney, Sydney, New South
Wales, Australia
Luca Montanarella
European Commission—DG JRC, Ispra, Varese, Italy
Inakwu O.A. Odeh
Faculty of Agriculture and Environment, The University of Sydney, Sydney, New South
Wales, Australia


Contributors

xi

Rajendra Prasad
Indian National Science Academy, and Division of Agronomy, Indian Agricultural Research
Institute, Pusa, New Delhi, India
Pedro A. Sanchez
The Earth Institute at Columbia University, Palisades, New York, USA
Yashbir S. Shivay
Division of Agronomy, Indian Agricultural Research Institute, Pusa, New Delhi, India
Bijay Singh
Department of Soil Science, Punjab Agricultural University, Ludhiana, Punjab, India
James A. Thompson
West Virginia University, Morgantown, West Virginia, USA
Zhi-Hong Xu
Environmental Futures Research Institute, Griffith University, Nathan, Queensland,

Australia
Bangbang Zhang
The College of Resources and Environmental Science, China Agricultural University,
Beijing, China
Gan-Lin Zhang
State Key Laboratory of Soil and Sustainable Agriculture, Institute of Soil Science, Chinese
Academy of Sciences, Nanjing, PR China
Qingpu Zhang
The College of Resources and Environmental Science, China Agricultural University,
Beijing, China


PREFACE
Volume 125 of Advances in Agronomy contains six cutting-edge reviews by
internationally recognized scientists. Chapter 1 is a state-of-the-art review
on the use of novel oxygen isotope ratios of phosphate to assess phosphorus
cycling in soil and water environments. Chapter 2 is a timely overview of
agronomic biofortification of cereal grains with iron and zinc. Chapter 3
presents exciting advances on the Global Soil Map, a digital soil map that
provides a fine-resolution global grid of soil functional properties.
Chapter 4 covers the effect of fertilizer intensification and its impacts in
China’s Huang Huai Hai plains. Chapter 5 discusses nutrient management
and use efficiency in South Asian wheat systems. Chapter 6 is a state-of-theart review on ammonia-oxidizing archaea and their important role in soil
acidification.
I am most grateful to the authors for their excellent contributions.
DONALD L. SPARKS
Newark, Delaware, USA

xiii

CHAPTER ONE

Advances in Using Oxygen Isotope
Ratios of Phosphate to
Understand Phosphorus Cycling
in the Environment
Deb P. Jaisi*,1, Ruth E. Blake†

*Department of Plant and Soil Sciences, University of Delaware, Newark, Delaware, USA
†
Department of Geology and Geophysics, Yale University, New Haven, Connecticut, USA
1
Corresponding author: e-mail address:

Contents
1. Introduction
1.1 Origin of phosphorus
1.2 Overview of P chemistry and P cycling
2. Stable Isotope Systematics: Oxygen Isotope Ratios of Phosphate
2.1 Oxygen isotope ratios of phosphate: Historical development
2.2 Apatite versus dissolved inorganic phosphate
2.3 Dissolved Pi–water oxygen isotopic fractionation and calibration
2.4 pH effect on Pi–water oxygen isotopic fractionation
2.5 Resistance to Pi–water O exchange and inorganic hydrolysis
2.6 Phosphate in the environment: recent developments
3. Organic Phosphorus and Isotope Effects of Organic Phosphate Mineralization:
Enzyme- and Substrate-Specific Isotope Effects
4. Measuremnt of Oxygen Isotope Ratios of Phosphate in Sediments, Soils, and
Natural Waters

4.1 Processing of dissolved phosphate in water for silver phosphate precipitation
4.2 Organic phosphorus and isotope effects of organic phosphate mineralization
4.3 Extraction of soil/sediment P and processing for silver phosphate precipitation
4.4 Methods of measuring oxygen isotope ratios in phosphate
5. Isotope Effects of Abiotic and Biotic Processes Involving Phosphates
5.1 Fractionation during abiotic processes of sorption, desorption, and mineral
transformation
5.2 Bioavailability and cycling of phosphate at the mineral-water interface
5.3 Fractionation during transport and mobilization of phosphate
5.4 Marine sediments with multiple pulses of authigenic phosphate precipitation
5.5 Detrital phosphate from different provenances

Advances in Agronomy, Volume 125
ISSN 0065-2113
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2014 Elsevier Inc.
All rights reserved.

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Deb P. Jaisi and Ruth E. Blake

6. Application of Oxygen Isotope Ratios in Phosphate to Understand P Cycling in
Soil Environments and Agricultre
7. Concluding Remarks and Perspectives
Acknowledgments
References

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Abstract
Phosphorus (P) is universally recognized as an essential nutrient for all known forms of
life and a key element in mediating between living and nonliving parts of the biosphere.
Here, we provide a comprehensive review of the development of oxygen isotope
methods of phosphate and application to understand the biogeochemical cycling of
P. With the advent of robust analytical techniques able to accurately determine stable
oxygen isotope ratios in phosphate (d18OP) and the increased understanding of isotope
effects from controlled process- or reaction-based studies, d18OP values have been
increasingly applied to identify sources and cycling of P in many natural environments.
Because different sources have distinct isotopic compositions and various processes
impart specific isotopic fractionation or produce distinct pathways of isotopic evolution,
application of d18OP values as a tracer for P in biogeochemical processes is expected to
continue to expand as an exciting field of research in the future.

1. INTRODUCTION
1.1. Origin of phosphorus
Phosphorus (P) in Greek mythology is “FosjόrοB” meaning “lightbearer.” The element P was first produced accidentally by a German physician, Hennig Brand (ca. 1630–1692), after distillation of evaporated urine
in the hope of changing metals in urine into gold. It is presumably the reduction of phosphate by pyrolytic carbon (Goldwhite, 1981) that produced elemental phosphorus. Early Christians noted the use of phosphorus as
“perpetual lamps” that glowed in the dark. The glow of phosphorus originates from chemiluminescence during aerial oxidation of elemental (white)
phosphorus. Similarly, ammonium sodium hydrogen phosphate
tetrahydrate (NaNH4HPO4Á4H2O) was historically used by alchemists as
“microcosmic salt.” Thus, the employment of P for useful purposes started
long ago in human civilization.
P is the eleventh most abundant element in the Earth’s crust with a crustal
abundance of 0.099%. It is widely distributed as orthophosphate ðPO4 3À Þ in
soils, rocks, oceans, all living beings, and in many man-made materials (e.g.,
pharmaceuticals, agrochemicals, food additives). However, the importance

3

Oxygen Isotope Studies of Phosphorus Cycling in Soils

of P as a nutrient was not realized until the mid-1800 s. Since its discovery as
a plant nutrient and its extraction from phosphorite rocks to produce fertilizers, other applications of P in military, medical, technological, and nutritional applications have greatly expanded in recent centuries.

1.2. Overview of P chemistry and P cycling
1.2.1 P chemistry
P has atomic number 15, atomic mass 30.97, and its electron configuration is
1s2 2s22p6 3s23p3. The promotional energy 3 s ! 3d orbital in P is small
enough to allow vacant d-orbitals to participate in bonding and forming
hybridized orbitals. This ready availability of d-orbitals permits a relatively
large number of potential configurations of electrons around the nucleus and
therefore accounts for the origin of diverse P compounds. Similarly, the
higher contribution of the d-orbital leads to an effectively large atom with
low electronegativity and greater polarizability (Corbridge, 1985). These
properties along with high first ionization energy (10.48 eV) result in overwhelmingly covalent character of P in chemical reactions. Its coordination
number varies from 1 (P0, elemental P) to 6 (PCl6 À , phosphorus
hexachloride), and its oxidation state from À3 (PH3, phosphine gas) to
þ5 (PO4 3À , phosphate) (Fig. 1.1). These properties are likely responsible
for the ubiquity of P-containing compounds in Earth environments
(Westheimer, 1987).

O

H
Inorganic P

H−P

P

O

H

–3
R
R −P

Organic P

R

Orthophosphate

Elemental P

Phosphine

Oxidation
states

O−P−O

–1
R
R− P=O
R

Trialkyl phosphine Phosphine oxides

0

+1
O
R − P = OH
R

+3
O

+5
O

R − P = OH

R −P−O −R

OH

O −R

Phosphonic acid
Phosphenic acid
Phosphate ester
OR

RO − P

OR
Phosphite ester

Figure 1.1 Oxidation states of P and examples of organic and inorganic compounds in
different oxidation states. In general, P compounds with low oxidation state are less
common on Earth.


4

Deb P. Jaisi and Ruth E. Blake

Inorganic orthophosphate (referred to as Pi hereafter), the most prevalent
form of P in the lithosphere and biosphere, is a compound in which the
P atom is surrounded tetrahedrally by four oxygen atoms (i.e., in þ5 oxidation state and 4 coordination number). A variety of condensed phosphates
including pyrophosphate and polyphosphate originates from sharing of oxygens in PO4 3À ions. The next most common form of P, organophosphorus
compounds, is substituted phosphate esters in which P and C are linked
through O as a PdOdC bond. Also common in biological system are phosphosulfur compounds such as APS (adenosine phosphosulfate) a key intermediate in bacterial respiration of sulfate ðSO4 2À Þ which is highly prominent
in marine sediments. Phosphonates, in which P(5þ) is bonded directly to C,
in PdC linkage, were once thought to be relatively rare and insignificant
forms of P in earth environments. Recent discoveries, however, have shown
the widespread occurrence of phosphonates throughout the world’s oceans
(Clark et al., 1999) and identified their role as P sources for primary oceanic
productivity (Dyhrman et al., 2006), sources of atmospheric methane (Karl
et al., 2008), and possible role in prebiotic earth chemistry (Glindemann
et al., 1999; Pasek, 2008). These developments have drawn new attention
to the reactions, origins, and biogeochemical cycling of phosphonates over
the full span of earth’s history.
Unlike other essential elements in living beings, P was classically viewed
as a redox-insensitive element, with phosphate (oxidation state 5þ) being

the only redox state naturally present in the environment. However, existence of other P phases (Fig. 1.1) in the environment has been increasingly
realized (see above) and the redox chemistry of P has been explored (e.g.,
Metcalf and Wolfe, 1998; Metcalf et al., 2012; Pasek and Block, 2009;
Schink and Friedrich, 2000). The reduction of P (5þ) to 3 þ, 1 þ, or 3 À
redox states occurs under extremely reducing conditions. Although oxidative
degradation of these reduced compounds was known, reductive pathways to
produce PO3, PO2, and PH3 in the environment remained elusive. Most
recently, organisms have been found to readily utilize reduced-P compounds
(e.g., phosphite, hypophosphite) as a P source (Metcalf and Wolfe, 1998) and
undergo dissimilatory oxidation of phosphite (PO3)3À in marine sediments
(Schink and Friedrich, 2000; Schink et al., 2002). Sources of reduced-P in
the environment include phosphite in corroding meteorites (Bryant et al.,
2009; Pasek and Lauretta, 2005), phosphites and phosphides produced from
lightning-reduced phosphate (Pasek and Block, 2009), phosphite in geothermal waters (Pech et al., 2009), and phosphine gas in soils and sediments
(Glindemann et al., 2005). Similarly, phosphonate comprises about 5% of


Oxygen Isotope Studies of Phosphorus Cycling in Soils

5

the soil organic P (Cade-Menun et al., 2002) and a similar fraction in marine
organic P (e.g., Dyhrman et al., 2009; Kolowith et al., 2001).

1.2.2 Conventional methods for understanding of P cycling in the
environment
P is universally recognized as an essential nutrient for all known forms of life.
General biochemical reactions that involve P at the molecular level have
been studied intensively and elucidated in great detail (e.g., Lassila et al.,
2011; Maloney et al., 1990; Torriani-Gorini et al., 1994). Recent advances

in molecular and atomic spectrometric techniques such as NMR, Raman,
and X-ray techniques have enabled a unique capability of P analysis to measure submicron scale processes and reactions in biological and geological
samples without introducing complex and invasive pretreatment steps.
For example, significant advances in our understanding of biogeochemical
cycling and speciation of P compounds have been made through the application of 31P NMR methods in marine systems (e.g., Benitez-Nelson et al.,
2004; Clark et al., 1998, 1999; Paytan et al., 2003; Sannigrahi and Ingall,
2005), terrestrial and agricultural soils (e.g., Cade-Menun, 2005; Fuentes
et al., 2012; McDowell et al., 2007, 2008; Turner and Leytem, 2004),
and aquatic systems (e.g., Cade-Menun et al., 2006; Nanny and Minear,
1997). Similarly, natural abundances of cosmogenic radionuclides 32P and
33
P or spiking radionuclides into natural systems have been used to quantify
dissolved and sedimentary/soil P species, uptake and cycling, and transformations (e.g., Achat et al., 2009; Bu¨hler et al., 2003; Friesen and Blair, 1988;
Frossard and Sinaj, 1997; Frossard et al., 2011). X-ray absorption spectroscopy such as XANES (X-ray absorption near edge structure) and EXAFS
(extended X-ray absorption fine structures) have been used to understand
the processes, reaction mechanisms, and compositions of P in the water column (e.g., Diaz et al., 2012), marine sediments (e.g., Brandes et al., 2007),
and agricultural soils and manures (e.g., Beauchemin et al., 2003; Seiter
et al., 2008). Yet, despite the crucial roles of P within living organisms
(e.g., ATP, phospholipids, and nucleic acids), long-term ecosystem productivity, and global climate regulation, very little is known or reported on the
biogeochemical cycles, microbiology, and geobiology of P in textbooks and
critical review papers. This contrasts sharply with other important bioelements (e.g., C, N, S, and Fe) about which much is known and much
has been written in textbooks (e.g., Banfield and Nealson, 1997;
Konhauser, 2007; Madigan et al., 2006).


6

Deb P. Jaisi and Ruth E. Blake

Pi, the primary form of P in nature, has been greatly underutilized as a

geochemical indicator of P cycling in natural systems. This is due in large
part to the chemical properties of P. This lack of use as a geochemical tool
has limited our understanding on the links between specific biological activity and chemical signatures, as well as links between Pi and other macronutrients like C, N, and S in natural environments. Nonetheless, understanding
the biogeochemical cycling of major nutrients and trace metals (such as Fe)
that may play a critical role in limiting productivity over time is essential for
elucidating the current and future effects of both natural and anthropogenically induced changes in nutrient composition and global change.
1.2.3 Environmental problems associated with P
Because of the low stoichiometric need for P among other major nutrients
(ca. 106C: 16N: 1P; Redfield, 1958) for organisms, small amounts of
P addition or changes could cause severe impacts on water quality in the
receiving catchments or groundwater aquifers. Increase in anthropogenic
sources of P in open waters has caused eutrophication in socioeconomically
and politically sensitive regions of the world such as the Chesapeake Bay and
Gulf of Mexico in the United States, the Baltic Sea in Europe, and the
Changjiang Estuary in China. Series of studies in Lake Washington, Lake
Erie, and Ashumet Pond (Cape Cod, MA) have found P as a limiting nutrient for eutrophication (Boyce et al., 1987; Correll, 1999; Edmondson, 1970;
Schindler, 1977; Schneider, 1997) as with a vast majority of surface waters
(Wetzel, 1983). Similarly, in a soil development chronosequence, Pi concentration is high at first due to weathering, thereby biologically available
nitrogen limits plant production. As the soil development continues, Pi is
occluded in secondary minerals, lost by erosion, or sequestered as organic
P (Po) resulting in P as the limiting nutrient in the soil (Vitousek et al., 1997).
It has been increasingly realized that P concentration trends measured in
surface and subsurface waters classically used to study uptake, release, and
cycling of Pi are not sufficient to address more pressing questions on
P cycling. For example, regulatory agencies such as the Ohio EPA, in
its report by the Ohio Lake Erie Phosphorus Task Force, have identified
“sediment–nutrient interaction as a critical factor” to understanding nutrient
movement (Ohio EPA, 2010). In fact, understanding of nutrient–sediment
interactions is not straightforward for P because of its chemical properties.
For example, the amount of dissolved Pi is several orders of magnitude

smaller than the pool of rapidly cycling Pi in soil and sediments (Frossard
et al., 1995), and both biotic and abiotic reactions occur side by side with


Oxygen Isotope Studies of Phosphorus Cycling in Soils

7

Figure 1.2 Different P pools in soil including species transfer and bioavailability
(Frossard et al., 2011).

dissolution and precipitation of Pi at time scales varying from a few seconds
to several years, and with redistribution of Pi from one phase/pool to
another (Fig. 1.2) (Bu¨nemann and Condron, 2007; Fardeau, 1996; Jaisi
et al., 2011). During these complex cycling and transformation processes,
P is always surrounded by four oxygen with no change in redox chemistry,
therefore limiting the use of chemical means to trace P. Therefore, new tools
that are sensitive to these processes could advance classical understanding of
transfer, transformation, and immobilization of P in different environments.

2. STABLE ISOTOPE SYSTEMATICS: OXYGEN ISOTOPE
RATIOS OF PHOSPHATE
Although P has 23 isotopes varying in mass from 24P to 46P, only one
( P) is stable. Among radioisotopes, 32P and 33P have half-lives of 14.26 and
25.34 days, respectively, while all others have half-lives less than 2.5 min.
Being the only stable isotope, 31P cannot be used to understand P systematics
in the environment in similar fashion to other nutrient elements—N, C, and
S, which have multiple stable isotopes for study. However, under Earth’s
surface conditions, P occurs primarily as orthophosphate (PO4 3À , or Pi).
This permits the use of stable isotope ratios of oxygen in orthophosphate

as a potential stable isotope tracer of P in the environment (Blake et al.,
31


8

Deb P. Jaisi and Ruth E. Blake

1997, 2001, 2005; Colman et al., 2005; McLaughlin et al., 2006; Paytan
et al., 2002). Oxygen isotope ratios in orthophosphate are commonly
expressed as d18Op values defined as:


Rsample
18
d Op ¼
À 1 1000
ð1:1Þ
Rstandard
where R denotes the ratio of the heavy to light oxygen isotope (18O/16O)
and Rsample and Rstandard refer to the ratio measured in sample and standard,
respectively. The isotopic abundance is measured against a reference standard and is reported in % relative to the VSMOW (Vienna standard mean
ocean water) standard.

2.1. Oxygen isotope ratios of phosphate: Historical
development
Measurement of d18Op values was prompted initially as a means to overcome the problem of the missing oxygen isotopic composition of water
in applications of the carbonate–water thermometry equation to paleoclimate studies (Epstein et al., 1953; Karhu and Epstein, 1986; Shemesh
et al., 1988). The hypothesis was that Pi and carbonate (CO3) equilibrated
with the same water would also be at the same temperature, thereby allowing measurement of d18O values of coeval carbonate and Pi in the same

sample and using the combined Pi–water and carbonate–water thermometry
equations to solve for temperature with d18Ow of water canceling out.
However, a slope similar to the carbonate–water equation was found for
the Pi–water thermometry equation (Longinelli, 1966), thus, precluding this
mathematical solution and dampening the hopes of this approach. Application of Pi d18Op values then turned primarily to terrestrial and marine paleoclimate studies (e.g., Ayliffe et al., 1992; Bryant et al., 1996; Fricke and
O’Neil, 1996; Iacumin et al., 1996; Joachimski et al., 2009; Kohn et al.,
1996; Kolodny et al., 1983; Le´cuyer et al., 1999; Longinelli, 1984; Luz
and Kolodny, 1985; Luz et al., 1984; Vennemann et al., 2001; Wenzel
et al., 2001) based on the Pi–water thermometry equation first defined by
Longinelli and Nuti (1973a):
À
Á
ð1:2Þ
t ¼ 111:4 À 4:3 d18 Op À d18 Ow
where d18Op and d18Ow are the oxygen isotopic compositions of Pi and
water in %, respectively, and t is the ambient temperature (in  C).


Oxygen Isotope Studies of Phosphorus Cycling in Soils

9

The Longinelli and Nuti (1973a) calibration of the temperature
dependence of Pi–water oxygen isotope fractionations is based on trace
Pi in carbonate shells of marine invertebrates and is also consistent with biogenic apatite minerals (e.g., bones, teeth ) from primarily marine fish
(Longinelli and Nuti, 1973b). This empirical equation is the founding equation of Pi O-isotope paleothermometry and also widely used in interpretation of dissolved and soil Pi O-isotope systematics. It has been reassessed by
several researchers in separate empirical studies using other sample types
(Kolodny et al., 1983; Le´cuyer et al., 1996) and slightly revised equations
have been proposed.
Kolodny et al. (1983):

À
Á
t ¼ 113:3 À 4:38 d18 Op À d18 Ow

ð1:3Þ

Le´cuyer et al. (1996):
À
Á
t ¼ 112:2ðÆ15:3Þ À 4:20ðÆ0:71Þ d18 Op À d18 Ow

ð1:4Þ

The slopes of these equations are within error of Longinelli and Nuti
(1973a). The equation derived by Kolodny et al. (1983) has also been used
as an alternative Pi paleothermometry equation (e.g., Domingo et al., 2009;
Kastner et al., 1984, 1990). The Longinelli and Nuti (1973a) equation has been
generally assumed to also apply to dissolved Pi as well and has been used extensively to reconstruct the paleotemperatures of ancient marine and terrestrial systems as well as modern environments (e.g., Elsbury et al., 2009; Goldhammer
et al., 2011a; Jaisi and Blake, 2010; McLaughlin et al., 2004, 2006).
There could, however, be issues using Longinelli and Nuti (1973a) equation in recent studies that utilize different analytical techniques, namely,
TC/EA used for pyrolysis, and inconsistencies in normalization to various
NBS standards (Puceat et al., 2010). A potential issue could be the effect
of final precipitate (either BiPO4 or Ag3PO4) used for isotope analyses analogous to the cation dependent isotope effect found to be present in silicates
(see Section 4.4.1). The offsets in different measurement methods such as
online high-temperature reduction of Ag3PO4 in a glassy carbon reactor
(e.g., pyrolysis in TC/EA) (Vennemann et al., 2002), heating of Ag3PO4
with graphite in silica tubes (e.g., O’Neil et al., 1994), or conventional
fluorination (Crowson et al., 1991; Longinelli, 1966), are all corrected to
fluorination methods (Vennemann et al., 2002). Detailed analyses of these
differences were recently described by Puceat et al. (2010). The revised thermometry equation proposed by these authors includes $2.2% correction to

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Deb P. Jaisi and Ruth E. Blake

Figure 1.3 Pi–water fractionation equations developed by Longinelli and Nuti (1973a),
Kolodny et al. (1983), and Puceat et al. 2010). (Puceat et al., 2010).

the Longinelli and Nuti (1973a,b) equation which they believe to be based
on normalization to NBS standards (Fig 1.3):
Puceat et al. (2010):
t ¼ 124:6ðÆ9:5Þ À 4:52ðÆ0:41Þ

À

d18 Op À d18 Ow

Á

ð1:5Þ

It is important to note that if Puceat et al. (2010) equation is used, all former
interpretations of equilibrium/near-equilibrium isotopic compositions are
incorrect because the new equilibrium is $2.2% heavier than Longinelli
and Nuti (1973a) calculated equilibrium values (Eq. 1.2). This means most
of the dissolved Pi measured so far are below equilibrium and the few cases
where heavier than equilibrium isotopic compositions are found are probably
at equilibrium. Further discussion the calibration equation and identification
of universal equilibrium is needed (e.g. Longinelli, 2013; Puceat et al., 2013).

One additional note is that normalizing the measured raw d18Op values to
problematic standards such as NBS120c could introduce additional complications for data presentation and comparisons. A more detailed review of
these issues and potential approaches to resolve these issues is essential but
is beyond the scope of this chapter.


Oxygen Isotope Studies of Phosphorus Cycling in Soils

11

2.2. Apatite versus dissolved inorganic phosphate
It is important to recognize that in the widely used Longinelli and Nuti
equation “phosphate” occurs as trace Pi mineral in carbonate shells that is
assumed (by others) to be bioapatite as with other empirical equations listed
above (Eqs. 1.3–1.5). A large number of studies of both marine and terrestrial paleoclimate and paleoenvironments have employed d18Op analysis of
bioapatite materials ranging from fish scales to dinosaur bones that date back
as far as Ordovician–Silurian (e.g., Wenzel et al., 2001) due to the higher
fidelity of preservation of apatite minerals compared with carbonates.
The reliance on empirical data to determine bioapatite–water fractionations at low temperatures most relevant to earth surface environments stems
from the well-documented resistance of Pi to abiotic chemical reaction and
isotopic exchange at low temperature (Longinelli and Nuti, 1973a;
Longinelli et al., 1976; Luz and Kolodny, 1985; Tudge, 1960) (see further
discussion in Section 2.4). It was recognized early on that at low near-surface
temperatures Pi–water exchange occurs only through biologically catalyzed
reactions, and it was widely asserted that such reactions took place within the
cells of living organisms and assumed to be due to the involvement of Pi in
the large number of enzyme-mediated reactions therein (e.g., Ayliffe et al.,
1992; Kolodny et al., 1983; Longinelli, 1984; Luz and Kolodny, 1985).
A series of controlled laboratory and field studies to understand the relationship between d18O values of drinking water or ingested (solid) Pi and d18Op
values of newly formed bioapatite values in fish and mammals provided

definitive evidence of the ability of living organisms to reset Pi d18Op values
and to promote Pi–water O-isotope exchange and presumed equilibrium
fractionation (Kolodny and Luz, 1988; Kolodny et al., 1983; Luz and
Kolodny, 1985). In most studies that followed these pioneering works by
Kolodny, the Longinelli and Nuti (1973a) equation has been used universally to interpret measured d18Op values of apatite and dissolved Pi and, thus,
have assumed identical fractionation behavior between dissolved Pi and
mineral bioapatite. As well, authigenic apatite precipitated in sediments
and soils has been assumed to have identical fractionation behavior as bioapatite. The key underlying assumptions in both cases are a similar mode of
precipitation, and no fractionation between dissolved Pi and solid/mineral
Pi (bioapatite or authigenic soil/sedimentary Pi minerals).

2.3. Dissolved Pi–water oxygen isotopic fractionation and
calibration
Blake et al. (1997, 1998a,b) attempted the first controlled laboratory studies
of dissolved Pi–water oxygen isotope fractionations at low temperature by


12

Deb P. Jaisi and Ruth E. Blake

employing enzymes produced by microorganisms (bacteria) as well as cellfree purified enzymes. In these experiments, bacteria were supplied with an
organophosphorus compound (RNA) as a sole nutrient source, which
required that the cells release extracellular phosphoenzymes to the medium
in order to break down the large organophosphorus molecules into smaller
pieces and finally release Pi, C, and N needed for their growth. Dissolved Pi
in spent microbial growth media was then collected by adding calcium carbonate, which promoted abiotic formation of apatite as a replacement of the
calcium carbonate, similar to formation of authigenic apatite in carbonaterich sediments. The d18Op values of these so-called microbial apatites
showed a strong relationship between Pi–water fractionation and temperature that was similar but shallower in slope (0.15) compared to the Longinelli
and Nuti (1973a) equation (0.20). It was later discovered that incomplete

Pi–water O-isotope exchange and nonequilibrium isotope effects of the
extracellular phosphoenzymes employed by the bacteria decreased the slope
(Liang and Blake, 2009). Nonetheless, the similarity in temperature dependence and slope of dissolved Pi–water temperature relations from microbial
growth media was an important first demonstration of the potential of dissolved Pi in natural waters to record ambient environmental temperatures as
well as Pi–water exchange. Because these results were obtained from
laboratory-controlled experiments without field validation, generalizing
observed fractionations of Pi as produced universally by all microbes in
any environmental system is not advised. Therefore, an outstanding but
unanswered question is whether there are microbe-specific isotope effects
or a general effect that could be represented by a universal equation. Current
evidence from laboratory cultures shows dominance of microbial
P metabolism by a inorganic pyrophosphatase (PPase) like equilibrium isotope effect as Pi is processed more and more by cells (see below). In the case
of Po degradation by microbial enzymes, a major source of Pi in natural
waters (e.g., marine water column, sediment porewaters) initially produced
Pi records isotopic signatures associated with enzymatic hydrolysis to
varying degrees depending on subsequent processing of Pi by cells and/or
cell-free enzymes (Colman et al., 2005; Goldhammer et al., 2011a;
Liang and Blake, 2009). However, more data are needed, especially for
dissolved Pi, to ascertain whether this is generally the case in the natural
environments.
The demonstration of true equilibrium Pi–water oxygen isotope exchange
was first presented by Blake et al. (1998a, 2005). This was achieved by conducting controlled laboratory experiments using cell-free enzymes to


Oxygen Isotope Studies of Phosphorus Cycling in Soils

13

catalyze dissolved Pi–water exchange on laboratory timescales below 30  C.
These researchers demonstrated the approach to steady-state Pi–water

O-isotope fractionations from opposite direction, and that the intracellular
enzyme PPase could account for both rapid, wholesale exchange of all four
oxygen atoms in Pi with water, and temperature dependence of equilibrium
Pi–water fractionations, consistent with previous empirical relations
observed for bioapatites, including the data of Longinelli and Nuti
(1973a). PPase is a ubiquitous intracellular enzyme that is present and highly
conserved across all three domains of life and recently dubbed “the great
equilibrator” (Blake et al., 2005). Although PPase can account fully for
observed equilibrium, temperature-dependent Pi–water O isotopic fractionation between apatite or dissolved Pi and water, this does not rule
out the possibility that other enzymes or combinations of enzymes can also
produce similar equilibrium isotope effects. The observed agreement of previous empirical bioapatite–water fractionations with experimentally demonstrated equilibrium Pi–water fractionations catalyzed by PPase suggests
that (i) Pi–water fractionations determined in previous empirical studies
were assumed correctly to be at equilibrium and (ii) the bulk of oxygen isotope exchange and temperature-dependent fractionation occurs between
dissolved Pi and water and not apatite and water.
Demonstration of temperature-dependent equilibrium Pi–water
exchange promoted by turnover (uptake and release) of dissolved Pi by
whole, intact microbial cells was presented by Blake et al. (1998b). In these
controlled growth experiments, Pi was supplied as the sole source of P for
growth and hexose sugar as the sole carbon source, which required the use of
the phosphotransferase phosophorylation pathway, which promotes the
“cycling” of Pi through the intracellular region of cells—that is, repeated
uptake of Pi into cells and subsequent release of Pi from cells back to the
growth medium. The observed equilibrium Pi–water exchange promoted
by intact cells was attributed to the action of intracellular PPase although
it does not exclude the possibility of equilibrium fractionation being promoted by other intracellular enzymes. Although normally located intracellularly, PPase could also exist outside of cells due to processes of cell lysis,
predation, and bacteriophage activity. Colman et al. (2005) invoked such
extracellular PPase activity as possibly contributing to the near-equilibrium
Pi–water O-isotopic fractionations measured by these researchers in the
marine water column and coastal estuaries.
At higher temperatures, above $70  C, dissolved Pi–water exchange can

be measured on laboratory timescales. Le´cuyer et al. (1999) performed


14

Deb P. Jaisi and Ruth E. Blake

controlled laboratory experiments to determine rates of abiotic Pi–water
exchange and equilibrium between 70 and 135  C. Their results produced
a new Pi–water thermometry equation as:
Lecuyer et al: ð1999Þ :
À 18
Á
ð1:6Þ
1
1
18
¼
3 d Op À d Ow þ 32:29ðÆ1:01Þ
T ð18:35ðÆ0:37Þ Â 10 Þ
This equation is analogous to other calibration equations (Eqs. 1.2–1.5),
the difference being T is the absolute temperature (in K). When
this equation was extrapolated to low temperature, it showed a significant
ca. þ8% offset from the Longinelli and Nuti (1973a) equation. These
authors explained this difference as due to kinetic isotope fractionation,
possibly due to preferential uptake of isotopically light P16O4 during
growth of organisms. This explanation, however, contradicted previous
studies on fractionation between carbonate and Pi where carbonate
was found to be 8.5–9.9% heavier than Pi (Bryant et al., 1996;
Iacumin et al., 1996; Zazzo et al., 2004). A similar offset value (9.5%)

has been shown to occur in apatite of young (Cenozoic) phosphorites
from the present sea floor (Shemesh et al., 1983). O’Neil et al. (2003)
later showed that the 8% offset observed by Le´cuyer et al. (1999)
was most likely resulted from the effects of pH on fractionations between
dissolved inorganic Pi species and water (see Section 2.4).

2.4. pH effect on Pi–water oxygen isotopic fractionation
All calibration equations for Pi–water fractionations mentioned above are
based on empirical measurements of mostly marine organisms (pH $8.2)
or mammals (pH $7.4) or conducted in the laboratory at near-neutral
pH. O’Neil et al. (2003) addressed the pH dependence of Pi–water
O-isotope fractionations from a series of abiotic experiments carried out
at relatively high temperatures (70–180  C). Pi–water fractionation was
found to be dependent on Pi speciation, which is controlled, in turn, by
pH. At low pH, the exchange between Pi and water is much faster because
of the greater presence of protonated species like H3PO4 and H2PO4 than at
high pH. Oxygen isotope fractionation between protonated species (at low
pH) and nonprotonated species (like PO4 3À ) (at high pH) was found to be
5–8%. This result is consistent with the analogous sulfate–water system
(Chiba et al., 1981; Hoering and Kennedy, 1957).


Oxygen Isotope Studies of Phosphorus Cycling in Soils

15

The pH dependence of Pi–water O-isotope fractionations has not been
investigated in natural systems or in controlled laboratory experiments with
microorganisms. Therefore, interpretation of d18Op values in high temperature and extreme pH environments requires calibrations of Pi–water
fractionations to be expanded and extended to more extreme conditions

of pH, temperature, and possibly ionic strength. However, as experimentally
validated and explained below (Section 2.5), no Pi–water exchange occurs
in the time frame used in routine laboratory processing of Pi samples at
extreme pH values.

2.5. Resistance to Pi–water O exchange and inorganic
hydrolysis
Phosphate oxygen is bound tightly to P such that in many Earth environments, inorganic exchange of oxygen between Pi and surrounding water
is essentially negligible even over geological time scales at low temperatures.
From the earliest studies known on oxygen isotopic exchange between Pi
and water including many other oxyanions (Winter et al., 1940), from equilibrium exchange calculations (Urey, 1947), and subsequent later studies
(Blake et al., 1997; Kolodny et al., 1983; Shemesh et al., 1983; Tudge,
1960), it has been consistently found that dissolved Pi does not undergo significant O-isotope exchange with water at low temperature. This is true
even at extreme pH values (10 M nitric acid and 14 M NH4OH) and
high-temperature (70  C) conditions used routinely to extract and purify
Pi from natural samples on laboratory time scales (Blake et al., 2010). Extrapolation of results of hydrothermal experiments performed by Le´cuyer et al.
(1999) to assess the temperature dependence of Pi–water oxygen isotope
exchange shows that the exchange rate is extremely slow at low temperatures (Fig. 1.4b). For example, it would take $6000 years for 10% of Pi oxygen to exchange with ambient water at 10  C. Solid-state mineral Pi such as
apatite, monazite, and other REE-phosphates are more resistant to
O-isotope exchange with water than dissolved Pi, even under prograde
metamorphic conditions (Donald et al., 2006) and have been suggested to
remain unaltered for an extremely long period of time, as far back as the
Archaean (Blake et al., 2010). A corollary of these observations suggested
by Blake et al. (2001) is that the evidence of low-temperature O-isotope
exchange between Pi and water, if present, should indicate the presence
or activity of life. This property of d18Op values in biotic and abiotic environments has made it a useful biomarker for the detection of “life activities”
in regions physically inaccessible such as deep subsurface, sub-seafloor, and


16

Deb P. Jaisi and Ruth E. Blake

A

B

10

0

Klebsiella aerogenes on
10 mM Pi + glucose, 25Њ C

−10
0

50
100
Time (in hours)

Extent of oxygen isotope
exchange

20

4

δ18OPO − δ18Owater

30
100%
80%
60%
40%
20%
0%
–4.0

–2.0

0.0
2.0
4.0
Time (in log yrs)

150 °C
100 °C

75 °C
30 °C

6.0

2 °C

Figure 1.4 Biotic and abiotic Pi–water isotope exchange (A) Rapid turnover and evolution of extracellular Pi–d18Op in bacterial P cycling, Klebsiella aerogenes at 25  C (Blake
et al., 2005). Note approach to equilibrium from opposite directions with Pi with different isotopic compositions; (B) Predictions of % Pi–water O-isotope exchange at different
temperatures based on Lécuyer et al. (1999).

extraterrestrial systems (Greenwood et al., 2003). Application of this concept
has largely been a founding basis for many recent studies aimed at understanding the biological cycling of P in several natural environments including marine environments, freshwater bodies, soils, sediments, and aerosols
(e.g., Angert et al., 2011; Elsbury et al., 2009; Goldhammer et al., 2011a;
Gross et al., 2013; Jaisi and Blake, 2010; Li et al., 2011; Tamburini
et al., 2012).

2.6. Phosphate in the environment: recent developments
In recent years, several aspects of the O-isotope systematics of Pi during
microbial metabolism and microbially mediated reactions in the marine
water column have been further characterized (Blake et al., 1997,
1998a,b, 2005; Colman et al., 2005; Goldhammer et al., 2011a;
McLaughlin et al., 2006; Paytan et al., 2002). Similarly, advances in methods
of sample purification and isotopic analysis, particularly a significant decrease
in sample mass requirement (Section 4.1), have led to rapid expansion of the
application of this tool to a wide range of sample types from soils, sediments,
and natural waters to biomass. A series of studies on the effects of predominant biogeochemical processes in the environment has applied Pi d18Op
values as a geochemical tracer. This includes basic understanding of isotope


Oxygen Isotope Studies of Phosphorus Cycling in Soils

17

effects associated with degradation of organophosphorus compounds—a
major source of P for microorganisms (Liang and Blake, 2006a, 2009), apatite precipitation (Liang and Blake, 2007), sorption, desorption, and mineral
transformation (Jaisi et al., 2010), transport (Jaisi, 2013), bioavailability of
specific P phases in sediment for microbial uptake (Blake et al., 2001; Jaisi
et al., 2011), and isotopic signatures associated with bacterial uptake of dissolved Pi (Blake et al., 2005). Other recent studies have targeted isotopic
signatures associated with specific P phases in soils and sediments (e.g., apatite) including development of methods to measure isotopic compositions of
extracted Pi (Angert et al., 2011; Goldhammer et al., 2011b; Gross et al.,

2013; Jaisi and Blake, 2010; Jaisi et al., 2011; Tamburini et al., 2010,
2012; Weiner et al., 2011; Zohar et al., 2010). These studies have increased
the realm of possibilities of applying d18Op values as a tool to understand
P dynamics in soils and sediments.

3. ORGANIC PHOSPHORUS AND ISOTOPE EFFECTS OF
ORGANIC PHOSPHATE MINERALIZATION: ENZYMEAND SUBSTRATE-SPECIFIC ISOTOPE EFFECTS
Orthophosphate concentrations in many aquatic systems are insufficient to meet the P demand of living organisms, thereby forcing them to
seek organophosphorous compounds (Po) (e.g., Benitez-Nelson, 2000).
Owing to large molecule size, Po has to be broken down extracellularly
by the action of freely dissolved or membrane-bound microbial/plant
enzymes. Many different types of extracellular enzymes (50 -nucleotidase,
peptidase, alkaline phosphatase, aminopeptidase, phosphodiesterase, chitinase, etc.) have been found in marine and terrestrial ecosystems. The composition and concentration of Po depends on type of biota, degree of
P demand, and nature of the environment (e.g., Ammerman and Azam,
1985; Dyhrman and Ruttenberg, 2006; Smucker and Kim, 1991). Intracellular phosphatase enzymes may also catalyze exchange or replacement of
O in Pi with water, however, including PPase, which is the only enzyme
known so far to catalyze wholesale equilibrium Pi–water oxygen isotope
exchange and temperature-dependent fractionations.
18
O-labeling techniques have long been used by biochemists to elucidate
the processes of oxidative phosphorylation. For example, Mildred Cohn
used 18O isotope and 31P NMR to probe phosphoenzyme reaction mechanisms (Cohn, 1953; Cohn and Hu, 1978). Most recently, electrospray ionization mass spectrometry (ESI-MS) has been used to detect and quantify


18

Deb P. Jaisi and Ruth E. Blake

18 16
each mass species present in PO4 (P18O16O3, P18O16

2 O2, P O3 O, and
P18O4 (Melby et al., 2011). These reaction mechanisms were of interest
to isotope geochemists because phosphoenzymes could catalyze Pi–water
exchange. Existence of such enzymes and enzyme-catalyzed exchange reactions was implicitly assumed to be responsible for the observed empirical
relations between mineral Pi in bones, teeth, shells, fish scales and the water
ingested by organisms—the same as ambient water in the case of fish and
other aquatic species, and drinking water or body water in the case of land
mammals (e.g., Longinelli and Nuti, 1973a,b; Kolodny et al., 1983, Luz and
Kolodny, 1985). The specific enzyme(s) responsible for this equilibrium
exchange, however, had not been investigated prior to studies of Blake
et al. (1998a, 2005). These authors found out that PPase catalyzed the
wholesale exchange of O (i.e., exchange of all four oxygen atoms in Pi with
oxygen in ambient water) between Pi and water that is required to achieve
O isotopic equilibrium between Pi and water (see Section 2.3).
Alkaline phosphatase (APase), on the other hand, is a non-specific
enzyme that hydrolyzes a variety of organic phosphomonoester compounds
as well as condensed inorganic phosphates such as pyrophosphate and even
phosphite (Metcalf and Wolfe, 1998), but unlike pyrophosphatase, the
APase reaction is unidirectional and thus does not promote wholesale
Pi–water exchange. Furthermore, it imparts a large kinetic isotope fractionation (À30 Æ 8%) during conversion of Po to Pi. The mechanism by which
a phosphohyrolase cleaves Po is schematically shown for monoester and
diester Po compounds in Fig. 1.5. For example, the O-isotope signature
of enzymatic Pi regeneration from Po can be expressed as (Liang and
Blake, 2006a, 2009):

For APase, 50 nucleotidase :

À
Á
d18 Op ¼ 0:75d18 OPo þ 0:25 d18 Ow þ F1

For PDase, RNase :

À
Á
À
Á
d18 Op ¼ 0:5d18 OPo þ 0:25 d18 Ow þ F1 þ 0:25 d18 Ow þ F2

ð1:7Þ
ð1:8Þ

where d18OP, d18OPo, and d18Ow are O-isotope composition of Pi
regenerated from nucleic acid, PO4 3À moiety groups bound to nucleic acids,
and ambient water O-isotope composition, respectively. F2 and F1 are fractionation factors associated with breaking the P-diester and P-monoester
bonds, respectively (corresponding to steps 1 and 2 in Fig. 1.5). It is important to note that these equations are valid for the enzymes/substrate studied


19

Oxygen Isotope Studies of Phosphorus Cycling in Soils

A

O

O
R2

O

P

O

R1

Phosphodiesterase

R2

O

O−
H

O−

P

+

HO

R1

O−

..

Phosphomonoester

O
H

B

O

O
R2

O

P
O−

..
H

O−

Phosphomonoesterase

R2

OH

+

O−

P

O−

O−
Phosphate

O
H

Figure 1.5 Hydrolysis of P-diesters. (A) Two-step bond cleavage reaction in diesters
includes the formation of P-monoester first with release of Pi and subsequent cleavage
of P-monoester bond and release of free Pi. Oxygen incorporated in Pi can also come
from an dOH on RNA structure in the case of RNase, other than H2O (not shown in
the drawing); (B) Second hydrolysis step breaks down monoester into alcohol and
orthophosphate (Liang and Blake, 2009).

by Liang and Blake (2009) but may not represent fractionations association
with hydrolysis catalyzed by other P-diesterases and P-monoesterases or
their specific substrate. More studies are needed to determine the representativeness of these specific results and their relevance to enzymes in natural
systems.
Results presented by Liang and Blake (2006a, 2009) suggest that the
O-isotope effect of APase and 50 nucleotidase monoesterase hydrolysis
depends on compound structure, but not on the chemical composition,
whereas hydrolysis of phosphodiesters DNA and RNA was compound specific. This promoted the development of structure-reaction-based models of
predicting isotope effects during cell-free enzyme as well as microbial degradation of Po. For the monoesters studied by these authors, one of the four
oxygen atoms in Pi comes from water (i.e., three out of four oxygen atoms
inherited from Po). The fractionation of incorporated and bulk water was

found to be different depending on the enzyme. For example, during 50 nucleotidase-catalyzed 50 -AMP hydrolysis, fractionation of water oxygen
(i.e., fractionation between incorporated water and ambient water) was
À10(Æ1)% but was $5% heavier for the APase-catalyzed hydrolysis of
the same 50 -AMP substrate.