Coupled Measurements of the D15n and the D18o of Nitrate As Tracers for Ocean Nitrogen

Coupled measurements of the d15N and the d18O 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

Preview

The goal of my proposed doctoral research is to study the behaviour of d18O 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 15N: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 d15N and d18O 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 1979).

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)

assimilation

NO2- (nitrite)

denitrification nitrification

(nitric oxide) NO

(?) (hydroxylamine)

(nitrous oxide) N2O NH2OH

N2 anabolism/ assimilation

nitrogen fixation NH4+ PON/DON

(ammonium) catabolism

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 a-, g-, 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.

Process / Tg N yr-1
sources
Pelagic N2 fixation / 110a - 330b
Benthic N2 fixation / 15c
River input / 25b - 76a
Atmospheric deposition / 30a
Total sources / 180 - 451
sinks
Water column denitrification / 80a
Sedimentary denitrification / 95a - 280d
Sedimentation / 25a
N2O loss / 4e
Total sinks / 204 - 389

a Gruber and Sarmiento (1997)

b Brandes and Devol (2003)

c Capone (1983)

d Middelburg et al. (1996)

e 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 15N. 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 because:

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 d-notation (in per mil units, ‰), as

d15N(‰) = - 1 x 1000 (1)

The standard is atmospheric N2, which in this notation has a d15N of 0‰. As illustrated in Figure 2, the d15N of NO3- increases progressively as nitrate is depleted from the culture medium by cellular uptake. The isotope effect e (also called fractionation factor) quantifies the relative magnitude of isotopic enrichment in the reactant pool. e is a function of the ratio of the reaction rates (k14 and k15) of the two isotopes,

e = (1 - k15/k14) x 1000 (2)

Experimentally, this value is calculated from the integrated expression of the progress of the reaction according to the following expression,

d15Nreactant = d15Ninitial - e{ln(f)} (3)

where f is the fraction of reactant remaining, d15Ninitial is the d15N of initial reactant N pool, and e 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, e is the negative slope of the linear relation of d15Nreactant (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 e (Figure 2), such that the instantaneous product is defined as

d15Ninstant = d15Nreactant - e (4)

It follows that the integral of this expression describes the d15N of the integrated product, namely that of total accumulated cell mass (see Mariotti et al. 1981),