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Coupled measurements of the 15N and the 18O of nitrate as
tracers for ocean nitrogen processes
Ph.D. Proposal
by Julie Granger
Department of Earth and Ocean Sciences
University of British Columbia
Presented on March 3, 2003
The goal of my proposed doctoral research is to study the behaviour of 18O in
nitrate for biologically-mediated transformations pertinent to the oceanic nitrogen cycle.
Laboratory experiments will be aimed at documenting the behaviour of 18O in individual
biological reactions of nitrate for cultures of microorganisms involved in oceanic
nitrogen cycling. More specifically, I will investigate isotopic fractionation of 18O/16O
(and 15N/14N) effected by nitrate assimilation by marine phytoplankton species.
Similarly, 18O/16O isotopic fractionation by dissimilatory nitrate reduction will be
examined in laboratory cultures of marine denitrifiers. Subsequently, studies of coupled
N:18O fractionation during nitrate assimilation by marine phytoplankton will be
extended to separating the isotope effect for cellular nitrate transport from that for nitrate
reduction, to elucidate the physiological mechanisms of isotope fractionation and its
controls. Finally, I will analyze depth profiles of 15N and 18O of nitrate from locations
in the Eastern Tropical North Pacific to determine the active processes that effected the
resultant isotopic profiles. The observations previously accrued in the laboratory on the
behaviour of nitrate 15N and 18O will serve as a point of reference from which to interpret
observed distributions of isotopically enriched nitrate in situ, and ultimately the cycling
of nitrate in the water column. Analysis of Eastern Tropical North Pacific isotopic
profiles will also provide an opportunity to test the robustness of coupled 15N:18O isotopic
ratios of nitrate as a tracer for biological nitrogen transformations. Coupled estimates of
nitrate 15N:18O isotopic ratios /may potentially offer a wealth of information whose
significance with regards to nitrogen cycling is, as of yet, undetermined.
The oceanic nitrogen cycle
Nitrogen is a major constituent of living mass and thus a chief determinant in
metabolism and growth of open ocean algae. Consequently, the distribution and mean
concentration of nitrate in the ocean affect the global fertility of the sea and its
consequent exchange of gases with the atmosphere. As such, nitrogen has been proposed
as a major driver of the atmospheric CO2 changes that characterize glacial/interglacial
cycles. Increased nitrate consumption in polar surface waters during the last glacial age
is hypothesized to have effected the apparent CO2 decrease (Francois et al. 1997).
Enhancement of low-latitude productivity due to increased nitrogen fixation at low
latitudes also figures as a plausible scenario to explain low CO2 concentrations during the
last glaciation (Falkowski 1997). Constraining the pools and fluxes of nitrogen in the
modern ocean, as well as understanding the mechanisms that underlie biological nitrogen
transformations, are thus paramount to expanding current knowledge of ocean
biogeochemistry. Ultimately, more intimate knowledge of the ocean's nitrogen cycle
may lead to insight into its relation to global climate change.
A schematic representation of the oceanic nitrogen cycle is presented in Figure 1.
Nitrate (NO3-), figured at the top of the diagram, is the most oxidized species of nitrogen.
Biological reduction of nitrate catalyses the loss of an oxygen atom, resulting in nitrite
(NO2-). This transformation is characteristic of two distinct biological reactions termed
assimilatory and dissimilatory nitrate reduction. The former refers to the assimilation of
nitrate by algae (and heterotrophic bacteria - Allen et al. 2002) for N-nutrition: Nitrate is
internalized at the cell surface and then reduced intracellularly to ammonia, via nitrite.
Ammonia then serves as the primary template for amino acid synthesis. Living mass thus
generated at the surface ocean is subject to consumption by grazers, or alternatively it
may senesce as a result of nutrient starvation or viral lysis. These processes engender
nitrogen release from grazed and senescent cells, as ammonia (or rather, ammonium, the
cationic form at seawater pH) or dissolved organic nitrogen (DON). DON can further be
catabolyzed by bacteria back to ammonium. Ammonium at the surface ocean, which
originates solely from consumption/decomposition of plankton, constitutes a choice
source of nitrogen for live phytoplankton. Primary production originating from the
utilization of ammonium as an N source is referred to as "regenerated production". "New
production," in contrast, is fuelled by nitrate freshly supplied to the surface ocean
(Dugdale and Goering 1967). Since, in a steady-state system, what enters the euphotic
zone (nitrate) must be exported back to depth (organic material), new production
measurements (e.g., 15N-labeled nitrate uptake rates measured for field sample
incubations) provide an estimate of total N export to the deep ocean (Eppley and Peterson
Deeper in the water column, ammonia released during organic matter
decomposition encounters a different fate. In the absence of light, nitrifying bacteria,
namely ammonia oxidizers and nitrite oxidizers, oxidize ammonia back to nitrate as a
means of securing reducing power to synthesize primary sugars from CO2. These
organisms do a distinctly thorough job of this, as no ammonium (or nitrite) is detectable
in deep water. Low concentrations of ammonium and nitrite do, however, accumulate at
the top of the nitracline and above in the euphotic zone, where multiple processes may be
operative simultaneously. At these depths, the supply of ammonium or nitrite may
exceed assimilation or oxidation rates. Phytoplankton cannot keep up with N supply as
light becomes progressively limiting with depth. Nitrifiers, on the other hand, may not be
able to use ammonium and nitrite fully because their activity is progressively suppressed
with increasing light levels.. Nonetheless, significant oxidation rates of ammonium and
nitrite are detectable at the nitracline and at shallower depths (Ward et al. 1989). So in
reality, nitrate is not only regenerated from ammonia below the nitracline, but also within
the surface mixed layer. This poses a caveat to the "new" vs. "regenerated production"
paradigm, which assumes no nitrate regeneration within the mixed layer. Ward et al.
(1989) report significant nitrate production within the mixed layer relative to nitrate
assimilation in the California current, implying that part of the nitrate assimilated is
functionally regenerated instead of new. Furthermore, the new production paradigm
assumes consumption of nitrate that is exclusive to photoautotrophs. Mounting evidence
reveals that a large fraction of nitrate is consumed by heterotrophic bacteria (Allen et al.
2002 and references therein), such that nitrate consumption cannot be equated with
carbon fixation. Euphotic zones throughout the oceans represent areas of dynamic N
cycling where operative N-processes yet remain poorly defined.
Dissimilatory nitrate reduction, the alternate pathway for biological nitrate
reduction, is also termed denitrification. In the absence of oxygen, denitrifying bacteria
use nitrate as a final electron acceptor to carry out respiration (reviewed in Zumft 1997).
Nitrite generated from this reaction can further be reduced sequentially to nitric oxide
(NO) gas, nitrous oxide (N2O) gas, and finally to dinitrogen (N2) gas (Figure 1)- whence
each intermediate serves as a terminal electron acceptor, albeit with sequentially
increasing redox potentials that provide for moderate to marginal electron gradients
within the respiratory chain.
The denitrification process is not widespread throughout the ocean, but occurs in
localized areas of high surface production and low oxygen source waters. The Arabian
Sea, the Eastern Tropical North Pacific, and the Peru Upwelling are known as major
areas of active water-column denitrification. Sediments underlying productive coastal
areas also pose as sites of substantial denitrifying activity (Table 1, Seitzinger 1988;
Devol 1991; Middelburg et al. 1996, Brandes and Devol 2003). Denitrification
represents the major sink for oceanic fixed nitrogen (Table 1). The magnitude of this loss
term is of utmost relevance for understanding the modern ocean nitrogen budget. Yet
due to the difficulty inherent in measuring and defining the extent of a process that is
variable in space and time, the loss of oceanic fixed N incurred from denitrification
remains poorly constrained (Codispoti et al. 2001, Brandes and Devol 2003).
NO3- (nitrate)
NO2- (nitrite)
(nitric oxide) NO
(nitrous oxide) N2O
nitrogen fixation
Figure 1. Schematic diagram of the processes and pools of N fundamental to the
cycling of N in the ocean. PON: particulate organic nitrogen. DON: dissolved
organic nitrogen. Nitrate is the most oxidized N species, while ammonium and
organic nitrogen comprise the most reduced species involved in the cycle.
Hydroxylamine is an intermediate species within the ammonia oxidation pathway
which does not accumulate extracelullarly. The dashed line designates a
physiological process that has been observed solely in vitro (Beaumont et al. 2002)
and whose oceanographic relevance is uncertain.
Denitrification in the ocean is countered by biological N-fixation, which involves
the catalytic reduction of dinitrogen gas to ammonia by nitrogen-fixing prokaryotes.
Much of the research on N-fixation in the marine environment has focused on the
cyanobacterium Trichodesmium. This genus inhabits low nutrient tropical and
subtropical seas where it often forms massive near-surface blooms of conspicuous
aggregate colonies (Carpenter and Capone 1992). Though Trichodesmium likely
contributes a significant fraction of total oceanic fixed nitrogen, a number of
cyanobacterial groups as well as -, -, and ß-proteobacteria are also potentially large
perpetrators of oceanic N-fixation (Zehr et al. 2001). Because N-fixation throughout the
ocean is spatially heterogenous, temporally stochastic, and thus, undersampled, the
generation of accurate estimates for global N-fixation rates has proven even more
challenging than for denitrification. Global N-fixation rates have been successively
revised upwards as more direct and indirect estimates are generated (Table 1, reviewed in
Karl et al. 2002), yet a recent model study by Brandes et al. (2003) suggests that even the
latest estimates may grossly underestimate marine nitrogen fixation rates.
The budget presented in Table 1 clearly illustrates that the sources and sinks of
fixed nitrogen to the ocean are presently poorly constrained, to the extent that it is not
even clear whether sources and sinks are in relative balance, or whether the ocean is
progressively losing or gaining fixed nitrogen.
Table 1. Fluxes for Sources and Sinks in the Global Marine Nitrogen Budget.
Tg N yr-1
Pelagic N2 fixation
110a - 330b
Benthic N2 fixation
River input
25b - 76a
Atmospheric deposition
Total sources
180 - 451
Water column denitrification
Sedimentary denitrification
95a - 280d
N2O loss
Total sinks
204 - 389
Gruber and Sarmiento (1997)
Brandes and Devol (2003)
Capone (1983)
Middelburg et al. (1996)
Nevison et al. (1995)
N isotopes as tracers of ocean N-processes
The Rayleigh model
The study of oceanic nitrogen cycling has been facilitated by the existence of a
stable isotope of nitrogen, namely
N. Naturally occurring nitrogen is comprised chiefly
of 14N, yet a minute fraction (0.36765 ± 0.00081 %) occurs as 15N, which possesses an
additional, stable neutron. The isotope generally has little effect on the chemical
properties of an element, as these are chiefly determined by electronic configuration. Yet
small differences in chemical behaviour of two isotopes of a given element do exist. For
a given element in fixed environmental surroundings, the kinetic energy (K) is constant.
Two isotopes of the same element have different masses but the same kinetic energy
K = 1/2mv2
such that masses of the same molecule (isotopomers) will have different velocities. An
example is water vapour. The lighter molecule has the higher velocity and can more
easily escape from the fluid phase. This causes isotopic fractionation, where the vapour
phase generated is relatively deplete in the heavier isotope, while the remaining fluid
phase is enriched with the heavier isotope.
The slight differences in nuclear mass between isotopes also affects the bond
energy, in that the bond strength of the heavier isotope is greater. In chemical reactions
that involve bond breakage, the energy barrier for the reaction of a molecule bearing a
heavier isotope is greater than that for the same molecule bestowed with the lighter
isotope. In biological reactions, mass-dependent differences in chemical behaviour often
result in isotopic fractionation, wherein molecules harbouring a lighter isotope (say 14N)
react more quickly than those that have 15N. As a consequence, throughout the course of
a biochemical reaction, the substrate being consumed becomes progressively enriched
with the heavier isotope, while the resultant product is relatively enriched with the lighter
isotope. This process is illustrated in Figure 2 for nitrate uptake by a marine diatom in
batch culture. On the y-axis, the isotope ratio of 15N to 14N is expressed in -notation (in
per mil units, ‰), as
15N(‰) =
15 14
N/ Nstandard
x 1000
The standard is atmospheric N2, which in this notation has a 15N of 0‰. As illustrated
in Figure 2, the 15N of NO3- increases progressively as nitrate is depleted from the
culture medium by cellular uptake. The isotope effect  (also called fractionation factor)
quantifies the relative magnitude of isotopic enrichment in the reactant pool.  is a
function of the ratio of the reaction rates (k14 and k15) of the two isotopes,
= (1 - k15/k14) x 1000
Experimentally, this value is calculated from the integrated expression of the progress of
the reaction according to the following expression,
15Nreactant = 15Ninitial - {ln(f)}
where f is the fraction of reactant remaining, 15Ninitial is the 15N of initial reactant N
pool, and  is the kinetic isotope effect of the transformation. The above equation
describes the Rayleigh model for isotope fractionation, which applies to reactions
occurring in a closed system (Mariotti et al. 1981). In practice, is the negative slope of
the linear relation of 15Nreactant (reactant = nitrate) vs. the natural logarithm of the
fraction of reactant remaining (f: nitrate/nitrateinitial).
As shown in Figure 2, total cell mass, i.e. the integrated product, also becomes
isotopically heavier throughout the reaction, since cells are consuming progressively
heavier nitrate throughout the course of the reaction. However, at any given moment, the
organic N being generated is always isotopically lighter than the reactant NO3- by a
difference of  (Figure 2), such that the instantaneous product is defined as
15Ninstant = 15Nreactant - 
It follows that the integral of this expression describes the 15N of the integrated product,
namely that of total accumulated cell mass (see Mariotti et al. 1981),
15Nintegrated = 15Ninitial + {ln(f)} x {f/(1-f)}
The Rayleigh model has been an invaluable tool to make sense of N-isotopic data
for processes occurring both in laboratory cultures, as well as in oceanic situations. An
alternative to the Rayleigh model is the steady-state model, in which reactant N is
continuously supplied and partially consumed, and residual reactant is exported at a
steady-state rate. However, this is beyond the scope of this overview.
Figure 2. Consumption of nitrate during growth of the marine diatom Thalassiosira
weissflogii in batch culture, and the concomitant increase in the 15N of nitrate (the
reactant). The estimated  for the integrated reaction is 11‰. Also plotted are the
calculated instantaneous 15N of growing cells (the instantaneous product) as well as the
calculated 15N of accumulated cells (the integrated product). Measurements of nitrate
and 15N of nitrate for the growing culture were measured by Granger and Sigman
Oceanic N-isotopic budget
Fractionation of N isotopes in the ocean reflects the biological processes active in
the water column. As such, N-isotopes have been used as a tool to elucidate N-cycling in
both the modern and paleo-ocean. The 15N of particulate organic nitrogen and of nitrate
in the water column show variations in magnitude that reflect the biological
transformations effected on ambient N. Similarly, organic-N residue stratified in deepsea sediment is also telling of past history of N-cycling and organic N sedimentation.
Figure 3 illustrates the current N stable isotope budget of the modern ocean. In
deep water resides the bulk of fixed nitrogen in the form of nitrate. Measurements of
deep ocean 15N throughout the seas converge on a relatively uniform value of 5‰
(Sigman et al. 2000). This value reflects the integrated signal of all localized N-isotopic
fractionation effected by major (biological) sources and sinks of fixed N in the ocean.
Nitrogen fixation, the dominant input term for oceanic fixed nitrogen (Table 1),
provides new nitrogen with a 15N of around -1 to 0‰, as measured in cellular N of Nfixer colonies collected at sea (Carpenter et al. 1997). By comparison, the 15N of
dissolved N2 is around 0.6‰ relative to atmospheric N2. Laboratory cultures of N-fixers
(Table 2) corroborate the apparent lack of N isotope fractionation associated with Nfixation, where fractionation factors () around 0‰ have also been measured.
Consequently, the 15N of organic material collected in shallow sediment traps in some
oligotrophic tropical gyres is relatively low (around 2‰), as particulate nitrogen sinking
out of the surface ocean bears the signature of N-fixation (reviewed in Karl et al. 2002).
The plankton mass that incorporates freshly fixed nitrogen and sinks out of the surface
ocean is decomposed and nitrified to nitrate that retains the low 15N imparted by Nfixation.
Were it not for the large isotope effect associated with denitrification (the
dominant sink for fixed nitrogen), the 15N of bulk nitrate in the deep ocean would
remain around 0‰. However localized pockets of denitrification throughout the ocean
impart a heavy 15N signal on resident nitrate. Three regions, the Peru Upwelling, the
Eastern Tropical North Pacific, and the Arabian Sea, account for most of global watercolumn denitrification in the ocean. A depth profile of nitrate concentration and 15N at a
location in Arabian Sea is plotted in Figure 4. Note the significant depletion of oxygen
that entrains denitrification. As nitrate (and nitrite) is used in lieu of oxygen to sustain
decomposition of organic material, the remaining nitrate pool becomes highly enriched in
N, in this case reaching upwards of 15‰ at the oxygen minimum (compared to 5‰ for
global deep ocean). Laboratory estimates of isotopic fractionation () by denitrifiers tend
to be high and variable (Table 2). Field values are similarly high, with more recent
estimates ranging between 20 to 30‰ (Table 3). The high degree of N isotope
discrimination associated with denitrification is thus reflected in the 15N of ambient
nitrate in denitrifying zones. On the whole, the magnitude and N isotopic signature of
denitrification, relative to the magnitude of oceanic N-fixation, amount to a global ocean
nitrate 15N of 5‰, as measured in deep water nitrate.
Neglected in the above simplification is the impact of sedimentary denitrification
on the global ocean N-isotope budget. The relative importance of sedimentary
denitrification as a sink for fixed N has been progressively revised upwards as our
Figure 4. Isotopic composition of nitrate (filled circles, open triangles, crosses) and nitrogen
gas (open circles) in the central Arabian Sea vs. depth. Open triangles represent 1993
hydrocast, filled and open circles 1994 data, and crosses 1995 data. All isotopic values are
per mil, vs. atmospheric Ne standard. Shaded region denotes depts with <10µM O 2
concentrations. Figure copied without permission from Brandes et al. (1998).
understanding of modern N cycle evolves (Table 1; Middelburg et al. 1996; Brandes and
Devol 2003). Unlike water-column denitrification, that in sediment is believed to impart
no isotope effect on nitrate because it is limited by the diffusion of nitrate to the
sediment. All nitrate supplied to the sediment is denitrified, such that no isotopicallyenriched nitrate pool remains (Figure 2; Brandes and Devol 1997). The global 15N of
mean ocean nitrate thus quantifies the net signal of fluxes, pools and respective isotope
effects for N-fixation relative to sedimentary and water-column denitrification.
The internal cycling of oceanic nitrogen, namely the cycle of nitrate uptake,
ammonification, and nitrification, has little effect on 15N of mean ocean nitrate. Nitrate
supplied from the deep ocean to the surface is completely consumed by resident plankton
in most of the global surface ocean. Although nitrate assimilation by phytoplankton is
associated with potentially large isotope effects (Table 2), complete nitrate consumption
pre-empts isotopically-enriched nitrate from remaining at the surface ocean (Altabet and
McCarthy 1985). Providing there are no alternate sources of fixed N to the surface (e.g.,
from N-fixation), organic material produced at the surface from nitrate originating from
deep water is imparted with the 15N of its source. And in a sisyphaean manner, the
organic nitrogen exported back to the deep ocean is remineralized to nitrate that has the
15N of deep ocean nitrate (Figures 1 & 3). Evidence of this process was presented by
Altabet (1988), who observed isotopic similarity between the annually integrated sinking
flux out of the Sargasso Sea mixed layer and thermocline nitrate from that region.
Sedimenting particulate nitrogen collected in sediment traps below the euphotic zone
showed a 15N identical to that of nitrate in the water underlying the euphotic zone,
showing close coupling between nitrate supply from deeper water to the surface and
sedimenting plankton mass.
Complete nitrate consumption at the surface ocean in part explains the relative
uniformity of deep water nitrate 15N (Sigman et al. 2000). The 15N measured for deep
ocean nitrate is relatively invariant both within and between deep ocean basins, estimated
around 4‰ in the North Atlantic to 6‰ in the North Pacific (Liu and Kaplan 1989, Liu et
al. 1996, Wu et al. 1997, Sigman et al. 1997, Sigman et al. 2000). Yet at high latitudes,
nitrate consumption by phytoplankton is not complete due to iron limitation of primary
production (e.g., Martin et al. 1994). Resident phytoplankton only consume a fraction of
the nitrate supply, such that a residual pool of 15N-enriched nitrate remains at the surface
ocean. Unlike denitrification, however, fractionation of surface nitrate from assimilation
does not effect any change in the global 15N budget of the ocean because N-isotopes are
merely redistributed in different water masses. Isotopically light fixed nitrogen is not lost
as N2, as is the case for denitrification. For example, Sigman et al. (2000) determined the
summer nitrate concentration at locations in the surface mixed layer of the Antarctic to be
around 25 µM, compared to 37 µM for source nitrate in the underlying water layer,
indicating incomplete nitrate consumption by phytoplankton. As expected, the 15N of
surface nitrate in the region was found to be enriched in 15N. In contrast, the 15N of
nitrate measured in the Upper Circumpolar Deep Water directly below was found to be
lower than 15N of nitrate at more northerly latitudes of the same water mass (where
nitrate use is greater). The diminished 15N of Upper Circumpolar Deep Water thus
appeared to reflect remineralization of isotopically light sinking organic N, the result of
incomplete nitrate use at the surface. Thus, surface processes resulted in a relatively
shallow and localized variation of nitrate 15N (Sigman et al. 2000). Such a variation
may be expected to be effaced later on, during deep winter mixing, and the nitrate 15N of
Upper Circumpolar Deep Water would then be restored to the value observed for the
northerly portion of the water mass. Incomplete nitrate consumption can result in local
variations in nitrate 15N that do not impact the relative homogeneity of deep ocean
nitrate. Globally, incomplete nitrate utilization incurs no net loss of fixed nitrogen from
the water column and thus no change in whole ocean 15N. Hence, processes proper to
the internal biological cycling of N do not act as a determinants of the global oceanic Nisotope budget.
Utility of N-isotopic measurements
In the modern ocean, studies of surface ocean nitrogen cycling, denitrification,
and nitrogen fixation have all used nitrogen isotopic patterns to investigate these
processes. Notably, Altabet and colleagues (Altabet and Deuser 1985, Altabet and
McCarthy 1985, Altabet and McCarthy 1986, Altabet 1988, Altabet 1989, Altabet et al.
1991, Voss et al. 1996) examined the linkages in the cycle of new nitrogen input into the
eupotic zone (in Atlantic oligotrophic warm-core rings, the Sargasso Sea, and the North
Atlantic), its utilization by phytoplankton, and transport to the deep sea, using 15N/14N
isotopic tracers. These studies pointed to a significant dichotomy between the 15N
signature of suspended particles collected in discrete water samples, compared to
sedimenting particles collected in sediment traps. A recurrent pattern observed in
oligotrophic gyres showed that sedimenting particles below the euphotic zone
consistently displayed a 15N identical to that of source nitrate (~3.5‰), while suspended
particles were generally light within the euphotic zone (`0.2‰), and heavier than source
nitrate below the photic zone. The similarity in 15N of sedimenting particles to source
nitrate implied that particles originated from nitrate supplied from below the euphotic
zone, which, in turn, implicated a negligible role of surface nitrogen fixation (in warmcore rings and in the Sargasso Sea) in supplying fixed nitrogen for plankton growth on
short (monthly) time scales. The low 15N value of suspended particles in the euphotic
zone (~0.2‰) was attributed to preferential export of 15N out of the surface layer by
sedimenting particles. And the transformation of isotopically light suspended particles
into isotopically heavy sedimenting particles was in turn taken as evidence of active
Table 2. Compilation of isotope effects () observed in laboratory cultures for N cycle processes.
Isotope effect ()
(NH4+ -> NO2-)
35 - 38‰
Nitrosomonas europaea a,b
Nitrosomonas marina b
Nitrosomonas C-113a b
Nitrosospira tenuis b
Nitrosomonas eutropha b
(NO3- -> N2)
20 - 30‰
13 - 20‰
Paracoccus denitrificansc
Pseudomonas stutzeri d
Pseudomonas denitrificans e
Nitrogen fixation
(N2 -> NH4+)
Trichodesmium sp. f
Azotobacter vinlandiie
NH4+ assimilation
15 - 19‰
Skeletonema costratum g
Mixed culture h
Emiliana huxlei (coastal) i
Emiliana huxlei (oceanic)i
Chaetoceros debilis i
NO3- assimilation
5 - 7‰
4 - 20‰
7 - 11‰
1 - 3‰
5 - 17‰
7 - 19‰
10 - 20‰
Thalassiosira pseudonana i, k
Emiliana huxlei (coastal) i
Emiliana huxlei (oceanic)i, k
Chaetoceros debilis i
Skeletonema costratum j
Isochrysis galbana j
Pavlova luteri j
Dunaliella tertiolecta j
Chroomonas salina j
Thalassiosira weissflogii j, k, m
Phaeodactylum tricornutum l
Thalassiosira oceanica k
Mariotti et al. (1981); bCasciotti et al. (submitted); cBarford et al. (1999); dWellman et al. (1968);
Delwiche and Steyne (1970); fCarpenter et al. (1997); hWaser et al. (1999); gPennock et al. (1996); iWaser
et al. (1998); jMontoya and McCarthy (1995); kGranger and Sigman (unpublished); lWada and Hattori
(1978); mNeedoba (unpublished).
Table 3. Isotope effects () from in situ estimates for N cycle processes.
Isotope effect ()
Chesapeake Baya
20 - 30‰; - 40‰*
Gulf of California h, i
25 - 29‰
Ammonia assimilation
Soil j
Trichodesmium colonies k
Western Tropical North Pacificl
6.5 - 8‰
Chesapeake Bay a
Nitrate assimilation
Eastern Tropical North Pacific b, c, d , e, f
Arabian Sea d, g
22 -27‰
Nitrogen fixation
Delaware Estuary m
5 - 20‰
Bacterial assemblage n
4 - 6‰
Southern Ocean o
Subarctic Pacific p
Equatorial Pacific q
Eastern North Pacific i
Horrigan et al. (1990); bVoss et al. (2001); dBrandes et al. (1998); eLiu and Kaplan (1989);*fCline and
Kaplan (1975); gNaqvi et al. (1998); hSigman et al. (submitted); iAltabet et al. (1999); jMariotti et al.
(1981); kCarpenter et al. (1997); lKarl et al. (1997); mCifuentes et al. (1989); nHoch et al. (1994); oSigman et
al. (1999a); pWu et al. (1997); qAltabet (2001).
processes such as macrozooplankton feeding and excretion of isotopically-heavy, fastsinking faecal pellets.
Montoya et al. (2002) reached opposite conclusions presenting concordant data
for the same oceanic region (Sargasso Sea) than those of Altabet (1988, 1989). In this
case, the low 15N of surface suspended particles was attributed to input of isotopically
light nitrogen from nitrogen fixation. Also emphasized is the putative role of
zooplankton grazing in creating sinking faecal material with a 15N somewhat higher
(~1‰ - modelled) than that for suspended organic material (~0‰ - measured). No
attempt is made to reconcile these conclusions with those of Altabet (1988, 1989).
However the interpretations put forth by Montoya et al. (2001) are in agreement with
observations of the occurrence of N-fixation at the surface of oligotophic gyres such as
the Sargasso Sea (Lipschultz and Owens 1996). Various factors point to a significant
role of nitrogen fixers in supplying a source of bioavailable nitrogen to the surface layer,
namely the actual presence of conspicuous colonies of nitrogen-fixers such as
Trichodesmium, numerous in situ measurements of N-fixation, as well as observations of
non-Redfield proportions between nitrate and phosphate at the base of the mixed layer
(Michaels et al. 1994, Gruber and Sarmiento 1997). Furthermore, N mass balance
calculations for the Sargasso sea identify a significant source of fixed nitrogen within the
main thermocline, which is inferred to be caused by nitrogen fixation (Michaels et al.
1996). Nevertheless, decisive interpretation of observed N-isotopic patterns in the
Sargasso Sea is still pending.
These studies also highlighted the important role of fast-sinking particles in
downward transport and redistribution of biochemical species in the deep ocean. A yet
unresolved caveat, however, is the observed decrease of sedimenting nitrogen 15N as a
function of depth during low flux periods; isotopic discrimination from bacterial
degradation would presumably result in the opposite trend. As for suspended particulate
nitrogen below the euphotic zone, its origin is ascribed to breakdown of larger, fastsinking particles. The 15N enrichment of these compared to sinking particles is
attributed to degradation processes. Based on measurements of bulk suspended and
sedimenting nitrogen, Altabet and colleagues computed local downward N fluxes, from
which they derived hypothetical N-supply rates required to achieve mass balance. As
such their work provided insight into internal cycling processes of N.
Of significance in the work of Altabet and others, as well as similar prior
observations by Miyake and Wada (1967) and Wada and Hattori (1978), is the recurring
observation of isotopically light suspended organic nitrogen at the base of the euphotic
zone, where nitrate is first discernible. This feature was interpreted as indicative of local
nitrate utilization, whence the existence of an inverse trend between particle 15N and the
relative fraction of nitrate utilized was established. Altabet and François (1994)
explicitly showed an inverse relationship between measurements of surface nitrate and
15N of near-surface particulate nitrogen along a north-south transect in the equatorial
Pacific. They also showed surface nitrate to be inversely related to core top sediment
15N. Application of sedimentary 15N as a paleotracer for surface nitrate utilization and
depletion had previously been the basis for interpretation of data from the Southern
Ocean (François et al. 1992, 1993) and the Mediterranean (Calvert et al. 1992). Altabet
and François (1994) thus provided important ground-truth for the application of
sedimentary 15N as a paleoceaographic tool. To date, studies of sedimentary 15N as a
paleo-recorder have shown glacial changes in both water column denitrification regions
and the Souther Ocean that would lead to reduced atmospheric pCO2 (Altabet et al 1995,
Ganeshram et al. 1995, François et al. 1997, Sigman et al. 1999b). Similar studies
conducted in the Equatorial Pacific infer a decrease in relative nitrate utilization in the
region during the last glacial maximum (Farrell et al. 1995, Altabet 2001).
Major determinants of the ultimate 15N signature of exported particulate N are
the 15N of the initial nitrate supply (15Ninitial), the isotope effect of nitrate uptake by
phytoplankton (), and the fraction of nitrate consumed (f). A number of studies provide
estimates of  for nitrate assimilation in different oceanic regions. These values are
derived from the Rayleigh model, using measurements of euphotic zone nitrate
concentrations, compared to the 15N of suspended or sedimenting particulate nitrogen,
or to the 15N of nitrate (Table 3). Most estimates converge on 5‰ throughout the ocean.
(Note that the similarity between global  estimates (5‰) and deep water nitrate 15N
(also 5‰) is incidental). Thus the isotope effect for nitrate assimilation appears relatively
invariant in the ocean. This contrasts isotope effects measured for laboratory cultures of
marine phytoplankton, which vary widely among and within species (Table 3). The
physiological factors that determine given isotope effects remain undefined, as the
mechanism of isotope fractionation during nitrate assimilation by phytoplankton is not
yet characterized. The  of 5‰ observed in the ocean likely reflects the mean of various
fractionation factors of resident plankton.
On local scales, N stable isotopic tracers can also provide information on the
source of ambient nitrate. Based on measurements of nitrate 15N in the Subantarctic,
Sigman et al. (1999) determined that Subantarctic surface water was supplied laterally
from Antarctic surface water, and not from the Subantarctic thermocline. Sigman et al.
(2000) also reported a 15N enrichment of Upper Circumpolar Deep Water of the
Southern ocean relative to underlying water masses, which they attributed to exchange
with low-latitude water carrying heavy nitrate from denitrification. In a similar example,
seawater off southern California was shown to be enriched in 15N, and isotopic maxima
were found to coincide with isopycnal levels of 15N-enriched water of the Eastern
Tropical North Pacific (Liu and Kaplan 1989). In contrast, Sigman et al. (2000) showed
relatively low nitrate 15N in the Subantarctic thermocline attributed to exchange with the
low-latitude thermocline, itself affected by mixing with low-nitrate surface water or by
the oxidation of newly-fixed N. Inputs of newly fixed N to the shallow thermocline of
the Pacific were also observed in particulate 15N measurements of sedimenting particles
by Liu et al. (1996) and Karl et al. (1997). Similarly, Brandes et al. (1998) documented a
pool of isotopically light nitrate in the thermocline waters of the central Arabian Sea,
which they attributed to local N-fixation.
N-isotopic fractionation associated with ocean denitrification has also been
investigated. In these studies, estimates of denit associated with water column
denitrification are derived from nitrate 15N measurements . The accuracy of field
estimates of denit are constrained by the difficulty of modelling isotope discrimination in
physically and chemically dynamic systems. This requires integration of terms such as
mixing and advection of water masses, as well as molecular diffusion. Thus estimates of
denit are limited by the capacity to define the physical environment and are therefore
subject to assumptions. Early estimates of denit varied considerably, ranging between
30‰ to 60‰ at different locations (in the Eastern Tropical North Pacific, the Peru
Upwelling, the Santa Barbara basin, and the Cariaco Trench) and depths (Cline 1973,
Cline and Kaplan 1975, Liu 1979). These values were derived from vertical or crossisopycnal diffusion models, consistent with the then current view of ventilation. Oxygen
minimum zones were seen as stagnant layers ventilated by mixing with surface and deep
oxygenated waters (Wyrtki 1962). Recent investigations of isotopic fractionation in
denitrifying zones derive denit using isopycnal oceanographic models that yield lower
estimates of denit, around 20 - 30‰ for locations in the Eastern Tropical North Pacific
and the Arabian Sea (Brandes et al. 1998; Naqvi et al. 1998; Voss et al. 2001). These fall
within the range of fractionation factors observed in lab culture studies (Table 3).
Less well understood are the localized impacts of dissolved organic nitrogen
(DON) release, ammonification, and ammonia and nitrite oxidation on the N-isotope
budget. DON likely constitutes the largest pool of fixed N after nitrate, yet its
concentration is rarely measured, and its isotopic composition has not been investigated.
Mounting evidence implicates DON fluxes as a significant component of N-cycling in the
surface ocean (Bronk and Ward 1999; Ward and Bronk 2001), and similarly DON
transformations may be important to the isotope dynamics of all N pools and consequent
implications ( al. 1991*; Karl et al. 1997*).
Ammonium is an intermediate in the regeneration of nitrogen that does not
accumulate in surface or deep ocean. It has hence been assumed to play a secondary role
in N-isotope dynamics. However the small pool that does accumulate at the base of the
euphotic zone ( ≤ 1 µM ) may prove significant to isotope flux dynamics because both
ammonia oxidation by nitrifiers and ammonium assimilation by competing
phytoplankton may impart large isotope fractionation to remaining pools (Figure 2;
Casciotti et al. submitted).
As seen above, N-isotopic tracers can provide qualitative and some quantitative
information on extant processes on local scales. However there have been few attempts
to use N isotopes in tandem with pool size and flux estimates to constrain the modern N
budget (Wada et al. 1975, Liu and Kaplan 1988, Altabet and Curry 1989). Notably, in a
recent exercise, Brandes and Devol (2003) constructed an isotopic mass balance of the
modern N budget based on current estimates of N pools and fluxes as well as isotopic
values for the various pools and processes. They concluded from their model that both
N-fixation and sedimentary denitrification were underestimated in current budgets, and
that these missing fluxes entailed a downward revision of the residence time of fixed N in
the ocean, from 3000 to 2000 years. This important study highlights the urgency of
constraining the magnitude of sources and sinks of fixed nitrogen in the ocean,
implicating a need to refine the various components of the ocean 's N-isotopic budget.
The 18O of nitrate as a tracer of biological N transformations
Cycling of oxygen within the N-cycle
Insights have been gained from the study of ocean N-isotope tracers, and
advances have been particularly significant in expanding knowledge of the paleo-ocean,
about which we know practically nothing. It has proven more challenging, however, for
N-isotopes to provide novel constrains for the modern ocean. In many cases, N-isotopes
have provided qualitative confirmation of operative processes, but have generally not
resulted in robust quantitative estimates of these processes. In cases where N-isotopes
have provided flux estimates, such as work by Altabet on internal N-cycling, the isotope
numbers have not corroborated numbers generated from alternate estimates. Though N
isotopes in these studies may point to missing components in understanding localized Ncycling, they do not resolve the discrepancies.
Yet the fate of N-isotope studies in the modern ocean is perhaps not be so bleak,
as an additional tracer for ocean N-processes has recently been proposed by Sigman and
colleagues (Sigman et al. 2001; Casciotti et al. 2002), which, when used in tandem with
N isotopes, promises to provide novel insights into the modern ocean N cycle. A recent
method pioneered by Sigman et al. (1999) involves the reduction of nitrate in seawater
samples to nitrous oxide gas by denitrifying bacteria, allowing for measurements of both
the 15N (Sigman et al. 1999) the 18O of nitrate (i.e., the 18O/16O ratio of nitrate Casciotti et al. 2002) at concentrations as low as 0.5 µM nitrate. The utility of this
additional tracer lies in the fact that measurements of the 18O of nitrate may act to
complement the processes in the oceanic N-cycle that are not fully captured by the
nitrogen isotopes.
A comparison of the plight of the 15N isotope throughout the N-cycle contrasted
to that of the 18O isotope clarifies the former statement (Figure 5). In the life of 15N, the
internal cycling of nitrogen, assimilation and remineralization, causes no change in the
isotopic composition of the nitrogen atoms involved (Figure 5). Net changes in whole
ocean (i.e., not local) 15N are driven solely by input and output processes, largely
dominated by nitrogen fixation and denitrification (Table 1, Figure 3). Nitrogen fixation
tends to lower the 15N of fixed nitrogen, while denitrification can cause an enrichment
of 15N. The trials and tribulations of 18O, while involved with fixed nitrogen, contrast
those of its nitrogenous partner (Figure 5). Oxygen is only transitory within the nitrogen
cycle, and its input and output processes are distinct from those of nitrogen. Nitrate
assimilation, organic matter decomposition, and subsequent nitrification do not represent
internal cycling processes, but rather comprise the input and output of oxygen to the Ncycle. In other words, oxygen comes aboard via ammonia and nitrite oxidation, and is
released back as water through nitrate and nitrite reduction. It follows that the 18O of
freshly generated nitrate (from nitrification) does not depend on the origin of the nitrogen
being nitrified, be it from newly fixed N, from denitrified N (as nitrite), from sedimenting
biomass of phytoplankton growing in nitrate-rich environments, or from sedimenting
biomass of plankton that completely assimilate the supply of surface nitrate. Because
oxygen is insensitive to these processes, its isotopic signature acts to complement that of
N, which bears the scars of previous N transformations.
Measurements of deep water nitrate suggest that the 18O of newly oxidized
nitrate is similar to that of seawater (Casciotti et. al. 2002, Figure 6). This may appear
surprising in light of the fact that biochemical studies of nitrifiers have established that
one of the three oxygens added to ammonia to form nitrate comes from dissolved oxygen,
while the remaining two originate from water (Andersson et al. 1983). One would hence
expect 18O values to partly reflect those of dissolved O2 in the ocean interior (which lie
between 23.8‰ to 35.5‰ relative to air - Bender 1990). However these same
biochemical studies also demonstrate a strong nitrifier-catalyzed oxygen exchange
between nitrite and water. Similarly, Casciotti et al. (2002) observe nitrite-to-water
oxygen isotope exchange in cultures of ammonia oxidizers, where less than one in six
oxygen atoms in the nitrite produced comes from oxygen. Oxygen exchange with water
is also plausible during nitrite oxidation, further replacing O2-oxygen atoms with H2Ooxygen atoms.
A depth profile of the 15N and 18O of nitrate at a location in the eastern
subarctic Pacific (the first such profile for ocean nitrate) shows that like the nitrogen atom
in nitrate, the oxygen atoms in nitrate are subject to isotopic fractionation, presumably
from plankton assimilation of nitrate at the surface (Casciotti et al. 2002, Figure 6). Of
further interest is the apparent similarity of the isotope effect observed for 15N compared
to that for 18O. Though seemingly counterintuitive, the fractionation factor associated
with the 15N of nitrate appears similar, if not identical, to that of the 18O of nitrate (ratio
of 1± 0.1 Casciotti et al. 2002). This pattern is corroborated by experiments that we
conducted with laboratory cultures of marine phytoplankton (Figure 7), where we
consistently observe a 1:1 ratio for O and N isotope discrimination, regardless of species,
culture conditions, or absolute isotope effect. Overall, preliminary observations suggest
that a 1:1 ratio for O and N isotope effects during nitrate assimilation may prevail
throughout the ocean.
Though 15N/18O coupling has not been investigated for laboratory cultures of
denitrifying bacteria, nitrate measurements in denitrifying zones of the Eastern Tropical
North Pacific suggest that the ratio for O and N isotope effects is also indistinguishable
from 1 (Figure 8, Sigman unpublished). However measurements of N/O isotope effects
Figure 5. Schematic diagram of the respective fates of N and O throughout N internal cycling. a)
Hypothetical example where nitrate at the surface ocean is completely consumed by resident plankton and
completely remineralized at the nitracline. Nitrate then advected to the surface has its original isotopic
siganture for both the N and O atoms. Note that neither the isotope effect associated with nitrate
assimilation nor that with ammonia oxidation imparts any change in the isotopic composition of the
respective products because the reactions go to completion. Also note the origin of the oxygen atoms from
water during ammonia oxidation. b) Only half of surface nitrate is consumed by resident plankton,
imparting a heavier isotopic signature for both N and O of nitrate, and lighter N isotopic mass for plankton.
Regeneration of sedimented plankton is complete. The isotopic signature of oxygen in nitrate resulting
from mixing of the remaining reactant pool to the regenerated pool differs significantly from that in the
previous example.
NO35‰ 0‰
 = 5‰
 = 15 - 35‰
5‰ 0‰
f = 0.5
 = 5‰
NO38.5‰ 3.5‰
NO35‰ 1.75‰
 = 15 - 35‰
1.5‰ 0‰
Figure 6. Nitrate15N and 18O (left) and nitrate concentration (right) for a depth profile
collected in July 1999 at Station P (50˚N, 145˚W) in the subarctic Pacific. Figure reproduced
from Cascoiotti et al. (2002).
 = 25‰
 N of NO 3 (‰, vs. starting value)
 = 1 ± 0.1
T. weiss.
T. oceanica
T. pseudo.
E. hux.
 = 5‰
ln(NO 3 /NO 3 initial )
 O of NO 3 (‰, vs. starting value)
Figure 7. Increase in 15N of nitrate vs. ln(f) for nitrate assimilation by marine phytoplankton
grown in batch culture (left). The fractionation factor (  is derived from the slope of the
relationship. Estimates of nitrate 15N vs. the corresponding 18O (right). Note the 1:1
relationshiop between 15N and 18O associated with nitrate assimilation. Data from Granger and
Sigman (unpublished).
for dissimilatory nitrate reduction in freshwater systems have suggested a 18O:15N ratio
of 0.5 to 0.7 for denitrification (refs). This discrepancy demands that the controls and
constraints on the 18O:15N ratio associated with isotopic discrimination during
denitrification be further investigated and characterized. If this ratio proves predictable
for biological N transformations in the ocean, it will provide a powerful tool to study
nitrogen cycle processes as well as patterns in ocean circulation.
Potential applications of the nitrate N and O isotopes in oceanography
Few data on the coupled 18O:15N of nitrate exist, and the existing ocean profiles
are not yet published and still wanting of original interpretation. Nonetheless, coupled
measurements of nitrate 15N and 18O may have the potential to provide novel insights
into biological N-cycling. One potential application is the use of coupled 15N/18O
measurements to determine the relative magnitude of assimilation to remineralization
within the surface mixed layer. Figure 5 illustrates two hypothetical scenarios of
euphotic zone production and remineralization cycles that result in distinct 18O: 15N
signatures. In first scenario (figure 5a), nitrate at the surface is completely consumed by
resident plankton and then completely remineralized at the top of the nitracline. Say the
15N of nitrate initially supplied to surface is 5‰, like that of deep water, and 18O is 0‰
(vs. standard mean ocean water), also like that of deep water. Plankton uptake of nitrate
has an isotope effect of 5‰, yet since all nitrate supplied is consumed (f = 1; equation 3),
there is no remaining nitrate pool that is isotopically enriched, and the plankton nitrogen
has a 15N of 5‰. The particulate nitrogen that sediments to the nitracline is ammonified
and completely re-oxidized within the mixed layer, resulting in new nitrate that has a
15N of 5‰ and a 18O of 0‰, and that resides within the mixed layer. Here again, the
undetermined magnitude of the isotopic fractionation imparted by ammonification as well
as the potentially large isotope effect associated with ammonia oxidation impart no
fractionation to the re-oxidized nitrate, as the reactant pools (the organic nitrogen and the
ammonia) are completely consumed (f = 1). Thus the 18O: 15N of the regenerated
nitrate is identical to that of the initial nitrate. So nitrate brought up to the surface again
has its original signature, that of deep ocean nitrate.
In the alternate scenario (Figure 5 b), nitrate supplied to the surface is not
completely consumed, such that a pool of nitrate enriched in both 15N and 18O remains at
the surface ocean. In this example, say half of the nitrate remains unconsumed and
fractionation of both N and O is assumed to follow a 1:1 trend. Using an isotope effect of
5‰ for nitrate assimilation, the remaining pool has a 15N of 8.5‰ and 18O of 3.5‰,
according to the Rayleigh model (equation 3). Particulate nitrogen has a 15N of 1.5‰
and has lost oxygen to water (equation 5). Say sedimenting organic nitrogen is
subsequently completely remineralized to nitrate at the top of the nitracline, within the
mixed layer (instead of below it), resulting in new nitrate that has a 15N of 1.5‰ and a
18O of 0‰ (from water). When this new nitrate pool is mixed back with the original
remaining pool, the 15N and 18O of the resulting nitrate are 5‰ for N and 1.75‰ for
oxygen. While the nitrogen retains its history throughout partial consumption and
remineralization, the oxygen in freshly oxidized nitrate is not linked to nitrogen during
remineralization and is thus insensitive to the source of N. Hence the 18O of nitrate
compared to 15N in this case provides a measure of the ratio of new vs. regenerated
nitrate in the euphotic zone. Incomplete consumption followed by remineralization acts
to raise the 18O: 15N ratio of nitrate in a water parcel. Such a signature in nitrate at the
ocean surface may thus be indicative of N regeneration occuring within the mixed layer,
such that traditional measurements of new production as plankton 15N-labeled nitrate
uptake would not equate export production (see discussion of new vs. regenerated
producion in "The oceanic nitrogen cycle" section).
Coupled 18O:15N measurements may also provide additional constraint on the
relative magnitude of denitrification compared to nitrogen fixation in a water parcel.
Estimates of denitrification can be extrapolated from the deficit in nitrate relative to
phosphate in the water column (e.g., Gruber and Sarmiento 1997), based on the
assumption that organic matter production and regeneration has a fixed N to P ratio (i.e.,
the Redfield ratio (Redfield et al. 1963)). However the potential impact of nitrogen
fixation on a water parcel that is being denitrified cannot be accounted for, as the two
processes cancel each other out. The same holds for 15N measurements, because
observed values reflect the difference between absolute denitrification rates (imparting a
large isotope effect) against any nitrogen fixation (with no isotope effect) that may have
occurred in a given water parcel. Extrapolation of denitrification rates based on Nisotopes alone may result in underestimates. Because all nitrate in deep water is
regenerated nitrate*, it follows that any enrichment in 18O of nitrate (relative to water at
0‰) below the euphotic zone is ostensibly the result of denitrification. Nitrogen fixation
does not impact the 18O of nitrate. Consequently, any positive deviation from a 1:1 ratio
of 18O: 15N in mid-depth waters (assuming denitrification does indeed fractionate N
and O with a 1:1 ratio) could be interpreted as indicating that the 15N of a given water
parcel bears the signal of nitrogen-fixation as well as denitrification (see below). From
there, the amount newly fixed N that has been added to the nitrate pool can be computed,
and "gross" denitrification rates revised accordingly. So the 18O of nitrate not only
provides constraint for quantification of two processes, but by same token, may provides
information on origin of nitrate in a water parcel.
*This does not take into account nitrate that remains unused at surface of Southern ocean.
Partial utilization results in isotope enrichment of both N and O. However, the fact that the
18O of deep nitrate is 0‰ implies that unused nitrate subducted to the deep ocean does not
effect a significant shift in the 18O of deep water nitrate.
Biochemical applications for N/O isotopes of nitrate
The mechanisms of isotopic fractionation during nitrate assimilation by marine
phytoplankton and during respiratory nitrate reduction by denitrifyers are not well
understood. Neither lab nor field measurements of isotope effects have provided robust
insight into what controls the magnitude of the isotope effect. Lab estimates vary widely
(Table2) and show no discernible pattern with respect to culture conditions, save an
increase in  at low light observed for a single diatom species (Wada and Hattori 1978,
Montoya and McCarthy 1995, Needoba unpublished). Field estimates of converge on
5‰ for nitrate assimilation, with some exceptions (e.g., Voss et al. 1996; Waser et al.
2000). As for denitrification, oceanic measurements hover around 25‰, with variability
that arises in part from the difficulty of modelling fractionation factor in a dynamic
As our present understanding of the fractionation mechanism for nitrate
assimilation is limited, interpretations of in vivo isotopic effects are mostly speculative.
The isotopic fractionation of N during nitrate assimilation is anecdotally attributed to the
catalysis of nitrate reduction by the enzyme nitrate reductase (NR). This hypothesis
stems from a number of factors. Measurement of N isotopic fractionation for purified
spinach nitrate reductase post a relatively high isotope effect of 25‰ (Ledgard et al.
1985). Nitrate reduction is believed to be the rate-limiting step in nitrate assimilation,
causing accumulation of an intracellular pool of 15N-erinched nitrate. Needoba and
Sigman (unpublished) have shown that a marine diatom does indeed accrue an
intracellular pool of isotopically enriched nitrate. Assuming fractionation is caused by
nitrate reductase, manifestation of the isotope effect in extracellular nitrate thence
requires that the organism be subject to significant rates of nitrate efflux. Shearer et al.
(1991) measured nitrate efflux rates and nitrate N-isotopic fractionation for a
cyanobacterium, and their observations were consistent with an NR based fractionation
step with nitrate efflux allowing for extracellular manifestation of the isotope effect.
Whether uptake or efflux steps during nitrate assimilation impart any isotopic
fractionation on nitrate is unknown. Mariotti et al. (1981) proposed that neither the
uptake nor efflux step have an intrinsic isotope effect on nitrate for assimilation by Pearl
Millet seedlings. Rather, they concluded that the isotopic fractionation observed during
nitrate assimilation is caused solely by NR. Whether fractionation occurs at the cell
membrane during nitrate influx or efflux has not been determined for eukaryotic
phytoplankton. The large variation in isotope effects observed both among and within
plankton species still remains unexplained and is mostly subject to conjecture (e.g.
Montoya and McCarthy 1995).
Coupled N/O isotope measurements may provide an important novel constraint to
resolve this quandary. As described above, equivalent isotope effects for N and O are
observed in experiments with marine phytoplankton cultures as well as in field data, and
this regardless of the amplitude of the isotope effects (Figure 7). Measurements of
isotope discrimination by phytoplankton NR in vitro may indicate whether coupling
occurs at the reductive step. Short term nitrate uptake experiments as well as pulse chase
experiments are also useful tools to determine efflux rates (Shearer et al. 1991) and the
fractionation imparted by uptake or efflux. Oxygen isotope measurements in this case can
provide an additional tracer to disentangle fractionation effected by the different steps
involved in nitrate assimilation (see below). A similar approach could be applied to
characterize the fractionation mechanism for dissimilatory nitrate reduction.
Proposed research
Coupled measurements of N and O isotope fractionation of nitrate offer a novel
and exciting avenue of research that promises to provide much insight into nitrogen
processes in the ocean. However, the behaviour of nitrate with respect to N and O
isotopes must clearly be understood before coupled 18O:15N variations in oceanic
nitrate can be used to disentangle operative N-processes in the water column, both
qualitatively and quantitatively. The goal of my doctoral thesis is thus to elucidate the
patterns and mechanisms of coupled N and O isotopic discrimination observed during
nitrate assimilation by phytoplankton. Through this work, I hope to provide a foundation
that will allow for pertinent interpretation of field observations of coupled N and O
isotope patterns in nitrate, and consequently lead to important qualitative and quantitative
constraints on oceanic N-cycling.
The proposed work consist of four sections. For the initial section, which is
nearly complete, I will measure N/O-isotope fractionation of nitrate by different marine
phytoplankton to discern patterns in the coupled isotope effects (Figure 7). For the
following part of my thesis, I will then repeat the former exercise with cultures of marine
and freshwater denitrifyers to characterize N/O isotope effects during dissimilatory
nitrate reduction. In the third part of the proposed work, I will attempt to uncover the
mechanism of isotope discrimination for nitrate assimilation by eukaryotic marine
phytoplankton (and perhaps for dissimilatory nitrate reduction by marine denitrifiers),
benefiting from added constraints provided by coupled nitrate N/O isotope measurements
to devise experiments aimed at deconstructing the fractionation mechanism. I also hope
to uncover the basis of the physiological controls on the fractionation mechanism by
determining environmental conditions that cause modulations in the magnitude of the
isotope effect for a given phytoplankton species. Finally, the last part of my work will be
based on a set of field measurements of nitrate 18O and 15N from the Eastern Tropical
North Pacific (measured by Sigman). I will interpret the data based on the patterns
observed in my lab studies on the behaviour of coupled N/O isotope discrimination. I
will generate models aimed at discerning the N-processes and their rates that corroborate
the observed N and O isotopic signatures. Each of the sections proposed above is
discussed in detail below.
Part 1: Isotope effects for N and O isotopic discrimination during nitrate assimilation by
marine phytoplankton.
I propose to study the change in isotopic composition of extracellular nitrate
during uptake by various marine phytoplankton. Both the nitrogen and oxygen isotopic
composition of nitrate will be quantified. These experiments will benefit from the advent
of a novel method to measure N and O isotopes of nitrate pioneered by Dr. Daniel
Sigman and colleagues at Princeton University (Sigman et al. 2001; Casciotti et al. 2002).
A number of methods exist to measure N isotopic composition of nitrate in seawater that
bare varying degrees of precision and sensitivity. The most effective method prior to that
of Sigman et al. (1999) entailed the reduction of nitrate to ammonia using DeVarda's
alloy, followed by diffusion of ammonia onto a basic glass-fiber filter encased between
teflon membranes (Sigman et al. 1997). The sensitivity of this method reached ≥ 5 µM
nitrate, which is still above many oceanographically relevant nitrate concentrations. The
more recently developed method by Sigman et al propounds greater sensitivity than preexisting methods for N-isotopic measurements ( ≥ 0.5 µM nitrate). Here, nitrate (and
nitrite) is reduced to nitrous oxide gas by denitrifying bacteria. The isotopic composition
of the gas is then measured by mass spectrometry. The analysis of nitrous oxide by mass
spectrometry not only provides the masses of the nitrogen atoms, but also those of
oxygen (Casciotti et al. 2003). The isotopic ratio of oxygen (18O to 16O) in nitrate was
not detectable with pre-existing methods. Coupled estimates of both N and O isotopic
ratios in nitrate are thus entirely novel and promising.
Dr. Daniel Sigman has agreed to measure the N and O isotopic composition of my
experimental nitrate samples, and will also be involved in the interpretation of the
accrued data. Dr. Sigman has will act as a co-advisor throughout the course of my
doctoral thesis. I believe his insights and experience will prove invaluable to my
progress, both intellectually and professionally.
To determine the isotope effect for N and O associated with nitrate assimilation
by marine phytoplankton, I will grow individual species in batch cultures and collect
liquid samples throughout exponential growth of the cells. Nitrate concentrations in the
samples will be measured and the samples will be sent to Dr. Sigman for isotope analysis.
Along with nitrate samples, subsamples of cell mass will be collected on glass-fiber
filters at each time point for isotope analysis of the particulate organic nitrogen.
Growth medium: Batch cultures of phytoplankton will be grown in the synthetic
seawater medium Aquil (Price et al. 1988/89) with varying concentrations of nitrate,
between 50 µM and 150 µM. Cells will be grown in acid-washed polycarbonate bottles
using trace-metal clean culture techniques. Metal concentrations will be buffered with
the metal chelator EDTA (ethylenediaminetetraacetic acid) following the procedure
outlined by Price et al. (1988/1989). The iron concentration will be manipulated in select
experiments as a means of modulating growth rates (i.e. imparting different
environmental conditions). Cells will be grown in a continuous-light incubator under
saturating light conditions. Growth will be monitored by quantifying cell densities with a
Coulter counter. The species of algae cultured will consist mostly of diatoms
(Thalassiosira weissflogii, Thalassiosira oceanica, and Thalassiosira pseudonana) as
well as a coccolithophorid (Emiliana huxlei).
Nitrate concentrations: Cells in the samples will be removed from spent medium
by gentle filtration onto a combusted GF/F, as intracellular pools of highly isotopically
enriched nitrate may become significant at high cell densities. Nitrate concentrations in
samples of spent medium will then be measured with an ozone-chemiluminescent NO
detector (NOx box) on-line with a reducing zinc vanadate solution (Garside 1982).
Particulate samples: Between 10 mL and 20 mL subsamples of the
phytoplankton cultures will be filtered onto a 25 mm pre-combusted glass-fiber filter
(GF/F), dried at 60˚C in a drying oven, and covered and pelleted in tin foil. Sample15N
will be measured using a VG PRISM mass spectrometer.
Nitrate N and O isotope analysis: As explained above, subsamples of spent
medium of actively growing phytoplankton cultures will be frozen at -20˚C and sent to
Dr. Sigman at Princeton University for isotopic analysis with the denitrifier method
(Sigman et al. 2001; Casciotti et al. 2003).
Data analysis: Isotope ratios of nitrogen and oxygen in nitrate, along with
concomitant nitrate concentrations will be fitted to the Rayleigh model to determine the
isotope effect associated with nitrate assimilation. Similarly, the isotope effect will also
be derived from the Rayleigh model from the N isotopic ratio of particulate organic
nitrogen samples. The patterns in isotope effect between the nitrogen vs. the oxygen
atom of nitrate will be compared, and the effect of growth conditions on the magnitude of
the N and O isotope effects will also be evaluated.
Part II: Isotope effect for N and O of nitrate associated with dissimilatory nitrate
The isotope effect imparted on both the nitrogen and oxygen atoms of nitrate by
denitrification will be investigated following a similar protocol than that used for the
phytoplankton cultures. Namely, subsamples of growing denitrifier cultures will be
collected as nitrate is being consumed. In many cases, nitrite accumulates in the growth
medium during dissimilatory nitrate reduction, such that a method to get rid of nitrite in
the subsamples will have to be devised in order to get at the isotope effect imparted on
nitrate alone. The range of species investigated will not be restricted to marine species,
but will also include freshwater isolates. As mentioned above, the pattern for coupled
N/O isotope fractionation associated with denitrification in freshwater profiles differs
from that observed in denitrifying oceanic regions. Comparison of freshwater and marine
denitrifying isolates may shed light on the causes of this pattern. Unlike the
phytoplankton experiments, particulate organic nitrogen will not be collected as most of
the cells are not retained by available glass-fiber filters (which are critical to N isotopic
analysis as they can be pre-combusted to remove trace N).
Growth medium (marine): Marine heterotrophic denitrifying bacterial isolates
will be cultured in Aquil (Price et al. 1988/89) modified for denitrifier growth (Granger
and Ward 2003). Cells will be inoculated in tri-laminate, gas-tight bags with a
polyethylene inner-lining. Oxygen in the medium will be consumed in the initial phase
of growth and nitrate concentrations will begin to decrease as oxygen becomes limiting.
Growth will be monitored from either the appearance of nitrite (measured
colorimetrically through azo-dye formation) or the disappearance of nitrate (measured as
above). Among experimental strains will figure Pseudomonas stutzeri.
Growth medium (freshwater): Freshwater denitrifiers will be grown in trilaminate gas-tight bags as above. Growth medium will consist of (microwave) sterilized,
filtered tap water supplemented with 10 µM phosphate, 50µM to 200µM nitrate, 0.5 gL-1
casein hydrolysate, 0.5 g/L-1 bactopeptone, f/2 vitamins, and 10 nM EDTA. Among the
strains to be cultured are Desulfovibrio desulfuricans (ask me why a sulphur reducer...),
and Paracoccus denitrificans
Removal of nitrite: The denitrifier method for measuring isotopic ratios of nitrate
concomitantly measures the isotope ratios of nitrite. Thus to distinguish isotope
discrimination on nitrate alone, nitrite must be removed from the spent medium. A
number of methods are available for the destruction of nitrite. However many of these
result in the formation of compounds toxic to the denitrifiers involved in nitrate isotope
analysis by the denitrifier method (Sigman et al. 2001), namely strains of Pseudomonas
chloraphis and Pseudomonas aerofaciens. For example, the reaction of nitrite with
sulfanilamide is an effective means of removing nitrite yet sulfanilamide is a potent
antibiotic (used to treat bladder infections). Toxicity caused by other available reactions
with nitrite is undetermined. I will thus test various nitrite removal methods and adopt
one that does not result toxicity to P. chloraphis and P. aerofaciens. Nitrite will thence
be removed from the culture samples, and the latter will be stored frozen until isotopic
Part III: The isotope fractionation mechanism during nitrate assimilation by marine
The isotope effect observed during nitrate assimilation by marine phytoplankton
may be imparted by one or more of the individual steps involved in nitrate assimilation.
The initial reaction involves the uptake of nitrate into the cell whence it joins an
intracellular pool of indeterminate size. Nitrate uptake is countered by nitrate efflux,
which commands a lesser rate relative to uptake in order to allow for net nitrate
consumption. Intracellular nitrate is thence subject to reduction by nitrate reductase in
the cytoplasm. The magnitude of this reaction depends both on the concentrations of
substrate and that of enzyme, as well as on the availability of reductant (NADH,
NADPH). In this model, fractionation may be imparted by uptake at the cell surface or
by efflux, though the role of these processes in determining observed isotope effects has
not been investigated. NR most likely fractionates intracellular nitrate (Legard et al.
1988), and this signal may be propagated into the extracellular medium if nitrate
accumulates intracellularly and if efflux rates are significant.
To determine the fractionation mechanism, I will need to quantify the relative
rates of influx, nitrate reduction, and efflux, and determine the isotope effects (if any)
associated with each of the aforementioned steps, which together amount to net nitrate
assimilation. The putative role of each of the above steps in shaping the observed isotope
effect will be investigated. Initially, the fractionation factor of cytoplasmic nitrate
reductase isolated from T.weissflogii. will be measured in vitro. The magnitude of
fractionation will be investigated, and whether it retains a 1:1 ratio between nitrogen and
oxygen atom of nitrate. I will then attempt to devise experiments to determine the
relative rates of influx and efflux of nitrate for actively growing T.weissflogii. This will
be done by using tungstate as an inhibitor of nitrate reductase and measuring changes in
net uptake as NR is incrementally inhibited by increasing concentrations of tungstate in
the growth medium. This procedure was employed by Shearer et al. (1991) to determine
relative influx/efflux rates of nitrate for Synechococcus R2. These experiments will also
allow me to determine any net fractionation imparted by combined influx/efflux
processes. To complement the above observations, estimate of nitrate influx/efflux rates,
I will conduct pulse-chase experiments with 15N-labeled nitrate. The cells will be grown
in 15N-lableled nitrate and resuspended in growth medium containing unlabeled nitrate to
monitor the appearance of 15N in extracellular medium over short time scales. Finally,
the size of the internal nitrate pool of T. weissflogii, as well as its isotopic compostition
(as pioneered by Joe Needoba for intracellular nitrate 15N measurements in diatom
species). The above experiments should provide me with estimates of influx and efflux
rates and associated isotopic effects, intracellular nitrate pool size and respective isotopic
signature, as well as net nitrate assimilation derived from measurements of cellular N and
growth. Collating all this information should result in a coherent model for the
mechanism of nitrate uptake and nitrate isotopic fractionation. The validity of the
proposed model will be tested by modifying growth conditions that should effect
predictable changes in the observed isotope effect. The above experiments will be
repeated for another species belonging to another phytoplankton group to determine
whether the fractionation mechanism of T. weissflogii may be the prevailing fractionation
mechanism among algal species.
I may attempt to characterize the fractionation mechanism for dissimilatory nitrate
reduction. This will require different approach as denitrifiers are endowed with three
types of nitrate reductases, a soluble assimilatory-type, and two dissimilatory-types
subdivided into the respiratory and the periplasmic nitrate reductases. The complexity of
the denitrifying mechanism, however, may prove too challenging to pursue in the time
allotted for the progression of my thesis.
Part IV: Application of coupled N:O isotope measurements in nitrate to uncover
operative N-processes in the Eastern Tropical North Pacific.
The study proposed above should add to current knowledge to provide a nearly
complete scheme of the behaviour of N and O isotopes of nitrate for each catalytic step
relevant to its ultimate 15N and 18O signature. The work I propose to conduct for my
doctoral thesis will complement and hopefully constrain the remaining uncharacterized
steps involved in coupled N and O fractionation of nitrate transformations. Empirical
observations of the unique signatures in N and O isotopes associated with both nitrate
assimilation and denitrification, as well as knowledge of the mechanism underlying
isotopic fractionation during nitrate assimilation, will provide us with a well-constrained
foundation from which to interpret patterns in nitrate 15N and 18O observed in situ.
My goal is thus to investigate 15N and 18O isotopes of nitrate in field samples
and uncover the processes that result in observed patterns. The Eastern Tropical North
Pacific is an area characterized by a dynamic and complex nitrogen cycle where coupled
N:O isotope patterns of nitrate promise to uncover the multiple processes at play. The
Eastern Tropical North Pacific is an area of active upwelling where productivity at the
surface is relatively high. Sedimentation and decomposition of the organic material
produced at the surface results in quasi complete consumption of ambient oxygen during
respiration, which triggers respiration of nitrogen oxides by denitrifying bacteria. The
loss of fixed nitrogen (nitrate and nitrite) as N2 gas is apparent at mid-depth waters where
oxygen concentrations are at a minimum . Calculation of the N* parameter (Michaels et
al. 1996, Gruber and Sarmiento 1997), which is based the loss of nitrate relative to
phosphate assuming Redfield stoichiometry of source nutrients, provides a quantitative
estimate of nitrate loss at a particular depth, and also acts as a tracer in calculations of
local or global N-budgets (Gruber and Sarmiento 1997). A depth profile of nitrate and
N* at a station in the ETNP is presented in Figure 8. Also plotted are the 15N and 18O
values for ambient nitrate, which were measured by Dr. Sigman (unpublished). The data
presented in Figure 8 constitute the basis of the last part of my thesis, wherefore I intend
to uncover the processes that led to the observed nitrate isotopic signatures.
The strongly negative N* throughout the water column posted in Figure 8 reflects
the loss of nitrate to denitrification (O2 ≤ 3µM from 175 m to the bottom of the profile).
Comparison of the nitrate concentration to the N* data (i.e., N*/(N* + [NO3-]) integrated
for the presented depth range suggests that ~25% of the nitrate has been lost to
denitrification. The N* minimum is reflected in a concomitant 15N and 18O maximum
between 200 to 400 m. This associated increase in 15N fitted to a closed-system
Rayleigh model yields an isotope effect of 25‰. Though the assumption of a closed
system is a doubtless oversimplification, the isotope effect derived from the Rayleigh
model is yet remarkably close to the mean isotope effect of 25‰ reported for recent field
estimates (Table 3), as well as most culture-based estimates (Table 2). This suggests that
most of the denitrification occurs in the water column; a significantly lower fractionation
factor would be uncharacteristic of water-column denitrification, and would thence imply
advection of water from the sediment surface, where denitrification shows no isotope
effect (Brandes and Devol 1997; Sigman et al. submitted). Though patterns observed for
18O compared to 15N seemingly fit relatively closely, the data show that the 15N
maximum is deeper that the 18O maximum. The difference between the enrichment in
15N relative to deep water and the concomitant enrichment of 18O relative to deep water
is plotted in Figure 8. Deeper water shows a 1:1 trend between 15N and 18O, which
amounts to a relative difference of zero. Above 350 m, the progressive enrichment in
18O does not parallel that in 15N; the isotopic enrichment for 18O is greater than that
for 15N, up to a depth of 100 m, above which the trend is reversed. Assuming that 15N
and 18O fractionate with a 1:1 ratio during denitrification, then a number of plausible
scenarios may explain the observed trend.
At the top of the profile, at 50 m depth, 15N is more enriched than 18O relative
to deep water. Regeneration of 15N-rich organic material sedimenting from the surface
likely to accounts for this. The 18O value is not 0, like that of water, implying that the
nitrate is has been denitrified (or consumed) since its generation from organic matter
remineralization. Yet the relative excess in 15N compared to 18O suggests that plankton at
the surface produce 15N-rich biomass, presumably because the nitrate upwelled to the
surface from mid-depths is itself enriched in 15N (from denitrification). Below 100 m, the
balance progressively shifts towards an excess of 18O relative to 15N - the effect of
organic matter remineralization is waning and overtaken by the undetermined process
that effects the higher 18O. This shift may be attributed to the oxidation of newly fixed
N nitrate elsewhere in the Pacific, with circulation distributing this new nitrate
throughout the Pacific thermocline. Figure 8 provides estimates of the newly fixed N
burden in the nitrate pool, about 15% of the total nitrate at 200 m (mean of two estimates,
see Fig. 8 caption). If this interpretation is accurate, then the 18O enrichment provides
the signature of "gross" denitrification, while that for 15N along with N* are indicative
of "net" denitrifcation equal to "gross" denitrification minus nitrogen fixation.
An alternative explanation for the observed pattern involves nitification following
partial denitrification. Suppose dissimilatory nitrate reduction to nitrite is initiated in a
water parcel whose initial 15N and 18O signatures are 5‰ and 0‰ (like deep water). If
5% of the nitrate pool is reduced to nitrite with a fractionation factor of 25‰, the
remaining nitrate pool has a 15N of 6.3‰, while that of the nitrite pool is -19.4‰.
Similarly, the 18O of oxygen in the remaining nitrate is 1.28‰ and nitrite -24.4‰
(assuming 1:1 fractionation). Say resident nitrite oxidizers (nitrifiers) take a fancy to the
newly generated nitrite and oxidize the whole of it back to nitrate. The 15N of the nitrate
pool resumes its orignial value of 5‰, while the 18O of nitrate now has a signature of 0.41‰ (resulting from its incorporation of a new oxygen atom from water with a 18O of
0‰). A study by Lipschultz et al. (1990) in the Peru Upwelling showed that
dissimilatory nitrate reduction and nitrite oxidation do occur simultaneously in the water
column, and at comparable rates - thence making the above scenario plausible.
In light of these different scenarios, any effect of nitrogen fixation needs to be
addressed in the context of a numerical model that accounts for all of the N-cycle
processes effecting the 18O and 15N signature of nitrate in the water column. This
model cannot be restricted to in situ biological processes of denitrification ,
decomposition and nitrification, but must also account for depth-specific advection of
nitrate, its magnitude and its biochemical history, and the effect of molecular diffusion of
nitrate in the water column, the sequence in time of biological reactions, etc...
This work will present an exciting challenge for me as I am not yet familiar with
water-column modellization. I therefore cannot outline how I propose to model the data
presented in Figure 8. I intend to initiate this modelling exercise with over-simplified
scenarios, and gradually add to its complexity throughout the course of my theses work in hopes of finally disentangling the true operative processes reflected in the ETNP
( N -  O) (‰ , relative to deep trends)
depth (m)
[NO 3 ]
 O
 N
from newly fixed N
8 10 12 14
 N or  O of NO 3 (‰ vs. air or SMOW )
[NO 3 ] (µM)
Figure 8. A profile of nitrate isotopes (left), concentration (right), and N* (far right) from near tip of Baja
California. Denitrification causes the maxima in the 18O and 15N and the minimum in N* (far right).
The enrichment in 15N and 18O from deep water fits a trend of ~1. Using this trend, we calculate the
decrease in the 15N due to nitrification of newly fixed (green, centre) and the corresponding fixed N input
(hatched bars). the fixed N input is estimated from 2 assumption s about whether denitrification precedes
or follows addition of fixed N, providing a lower and upper limit to the estimate. In reality, the processes
may be simultaneous and circulation mixes their effects, such that an accurate estimate of the new N input
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