Nitrogen Isotopes in the Ocean

Nitrogen Isotopes in the Ocean☆ DM Sigman, Princeton University, Princeton, NJ, United States F Fripiat, Max Planck Institute for Chemistry, Mainz, Germany © 2019 Elsevier Ltd. All rights reserved. Introduction Terms and Units Measurements Models Processes Inputs Outputs Internal Cycling Nitrogen assimilation Remineralization Nitrogen Reservoirs Dissolved Nitrogen Nitrate Nitrite Ammonium Dissolved organic nitrogen Particulate Nitrogen Suspended particles Sinking particles Nitrogen in sediments and the sedimentary record Concluding Remarks Further Reading 263 263 264 264 266 266 266 268 268 270 270 270 270 274 274 274 275 275 276 276 278 278 Introduction Nitrogen (N) has two stable isotopes, 14N and 15N (atomic masses of 14 and 15, respectively). 14N is the more abundant of the two, comprising 99.63  0.02% of the N found on Earth. Physical, chemical, and biological processes discriminate between the two isotopes, leading to subtle but measurable differences in the ratio of 15N to 14N among the forms of N found in the marine environment, making the N isotopes a powerful tracer of oceanic processes. Nitrogen is a central component of marine biomass and one of the major nutrients required by all phytoplankton. In this sense, biologically available (or “fixed”, i.e., non-N2) N is representative of the fundamental patterns of biogeochemical cycling in the ocean. However, fixed N differs from other ocean nutrients in that its sources and sinks are dominantly internal to the ocean and biological, with marine N2 fixation supplying much of the fixed N in the ocean and marine denitrification removing it. The N isotopes provide a means of studying both the input/output budget (i.e., absolute sources and sinks of marine fixed N) and the cycling of fixed N within the ocean. In this overview, we outline the isotope systematics of N cycle processes and their impacts on the isotopic composition of the major N reservoirs in the ocean. For reasons of length and scope, we focus on N nutrients and organic N, avoiding treatment of the isotope dynamics of the dissolved gases (e.g., N2 and N2O). The information presented here provides a starting point for considering the wide range of questions in ocean sciences to which the N isotopes can be applied. Terms and Units Measurements of the ratio of the N isotopes are typically reported relative to atmospheric N2 (“Air”), which has a 15N/14N ratio of 0.36765%  0.00041%. Natural samples exhibit small deviations from this ratio, which are expressed in d notation (in units of per mille, %, vs. Air):
! 15 N=14 N Sample 15
d N ð% vs: AirÞ ¼ 1
1000 (1) 15 N=14 N Air In this notation, the d15N of atmospheric N2 is 0%. ☆ Change History: September 2018. D.M. Sigman and F. Fripiat updated the text and further readings, figures, and tables. This is an update of D.M. Sigman, K.L. Karsh, K.L. Casciotti, Nitrogen Isotopes in the Ocean, Encyclopedia of Ocean Sciences (2nd Edn), edited by John H. Steele, Academic Press, 2009, pp. 40–54. Encyclopedia of Ocean Sciences, 3rd Edition https://doi.org/10.1016/B978-0-12-409548-9.11605-7 263 264 Ocean Process Tracers | Nitrogen Isotopes in the Ocean “Isotopic fractionation” refers to the isotope ratio differences caused by discrimination between the different isotopic forms of chemical species in reactions and other processes. A distinction is often drawn between “equilibrium” and “kinetic” fractionation, the latter nominally involving unidirectional reactions. Nitrogen isotope variations in the ocean are typically thought of as dominated by kinetic fractionation associated with the biochemical conversions of N from one form to another. However, equilibrium fractionation is important in some cases, for example, in the protonation and deprotonation of ammonia (NH3) and ammonium (NHþ 4 ). Moreover, kinetic and equilibrium fractionation are related. For example, equilibrium fractionation can be framed in terms of the kinetic fractionations of two opposing unidirectional reactions. As with isotopic composition, special terms are used to characterize the amplitude of isotopic discrimination. The kinetic isotope effect, e, of a given reaction is defined by the ratio of rates with which the two N isotopes are converted from reactant to product: ! 15 k 15 e ð%Þ ¼ 1 14
1000 (2) k where 14k and 15k are the rate coefficients of the reaction for 14N- and 15N-containing reactant, respectively. For e  1000%, e is approximated well by the difference in d15N between the reactant and its instantaneous product. That is, if a reaction has an e of 5%, then the d15N of the product N generated at any given time will be 5% lower than the d15N of the reactant N at that time. Measurements The isotopic analysis of N has long relied on gas-source, isotope ratio mass spectrometry, which requires the conversion of the sample to a stable gas analyte. For decades, online combustion to N2 has been the standard method for the preparation of a N sample for isotopic analysis. With “off-the-shelf” technology, a typical sample size requirement is 1–2 mmol N per analysis, with analysis precisions near 0.1% (1 standard deviation). Ongoing innovation with such systems is allowing for much reduced sample sizes; the performance of these more sensitive systems will become apparent as more application studies are published. Gas chromatography followed by micro-combustion to N2 is improving as a technique for specific organic compounds, amino acids in particular, although the polarity of many N compounds requires derivatization, impacting precision. Liquid chromatography prior to micro-combustion is also advancing. There are standard methods of collection for most bulk forms of particulate N (PN) in the ocean. Shallow and deep samples of suspended PN are filtered onto glass fiber filters (or, more recently, membrane filters from which the sample can subsequently be extracted). Sinking PN is most often collected by sediment traps. Zooplankton or large phytoplankton (e.g., Trichodesmium colonies) can be picked from filtered samples or net tows. Particulates can be separated into size classes. Recent N isotopic studies have used flow cytometry to sort cells based on taxonomic characteristics and to separate photosynthetic from non-photosynthetic cells and living cells from detrital organic matter. In the case of dissolved forms of N, the species of interest must be converted to a form, typically a gas, that can be extracted from the seawater. Since the 1970s, the d15N of marine nitrate (NO3 ), nitrite (NO2 ), and ammonium (NHþ 4 ) have been analyzed by conversion to ammonia gas and collection of the cationic ammonium form in acid for subsequent conversion to N2 (often referred to as the ammonia “distillation” and “diffusion” methods). Recently, much more sensitive isotope analysis methods have been developed for nitrate and nitrite in which these species are converted to the gas nitrous oxide (N2O; e.g., the “bacterial” or “denitrifier” method and the “chemical” or “azide” method). The N2O produced by these methods is analyzed by a purge and trap system, followed by gas chromatography and isotope ratio mass spectrometry. One benefit is the ability to exploit the high sensitivity of N2O isotope analysis, due to the suitability of N2O for cryogenic capture and release and its low atmospheric background; as a result, these methods require as little as 2 nmol N per analysis. The N2O-based methods also allow for oxygen isotope analysis of nitrate and nitrite, a measurement not previously possible in seawater. For samples that are to be analyzed for nitrate isotopic composition, pretreatment methods to remove nitrite have been developed. The methods for conversion of nitrate and nitrite to N2O also provide an alternative route for isotopic analysis of other forms of N, so long as they can be converted to nitrate or nitrite (e.g., by peroxydisulfate or hypobromite oxidation). The most pertinent benefit of this approach is its high sensitivity, which has reduced the quantity of N necessary for analysis and thus has expanded the range of sample types accessible to isotopic analysis. It has been applied to a broad range of N forms and sample types, including: dissolved organic N (DON) and ammonium; flow cytometrically sorted particles from the upper ocean; and the organic N trapped within the mineral matrix of diatom frustules, foraminifera shells, coral skeleton, and fish otoliths. Models Two simple models, “Rayleigh” and “steady-state”, are frequently used to interpret N isotope data from the ocean. In both of these models, the degree of consumption of the reactant N pool (f ) is a key variable, and the d15N of the initial reactant N pool (d15Ninitial) and the kinetic isotope effect (e) are the two central isotopic parameters. If a unidirectional transformation proceeds with a constant kinetic isotope effect and if the reactant N pool is neither replenished nor lost from the system during the progress of the transformation, then the process can be described in terms of Rayleigh fractionation kinetics, which define the isotopic variation Ocean Process Tracers | Nitrogen Isotopes in the Ocean 265 of the reactant N pool (d15Nreactant; Eq. 3), the instantaneously generated product N (d15Ninstantaneous; Eq. 4), and the accumulated product N pool (d15Naccumulated; Eq. 5) as a given reservoir of reactant N is consumed (Fig. 1): d15 Nreactant ¼ d15 Ninitial e
f ln ðf Þg d15 Ninstantaneous ¼ d15 Nreactant d15 Naccumulated ¼ d15 Ninitial þ e
ff =ð1 15 (3) e (4) f Þg
ln ðf Þ (5) 15 where f is the fraction of the reactant remaining, d Ninitial is the d N of the initial reactant N pool, and e is the kinetic isotope effect of the transformation. These equations are simplified, approximate forms of the full expressions. They are typically adequate, but their error is greater for higher consumption (lower f ) and higher e. The Rayleigh model is often used to describe events in the ocean. For example, the Rayleigh model has been used to simulate the uptake of nitrate by phytoplankton during the spring-to-summer, high-productivity period in a sunlit surface layer that has become isolated from the subsurface by density stratification. The two extrema in f warrant noting. At a very low degree of reactant consumption (f close to 1), the d15N of the reactant is approximately that of the supply, and the d15N values of the instantaneous and accumulated products are similar and approximately equal to the d15N of the reactant supply minus the kinetic isotope effect. At a very high degree of consumption (f close to 0), the d15N of the reactant is extremely elevated. More importantly, at f close to 0, the d15N of the accumulated product approximates the d15N of the initial reactant supply; this reflects the need for isotope mass balance, with the entire reactant pool having been converted to product. The end-member alternative to the Rayleigh model is the steady-state model, in which reactant N is continuously supplied and partially consumed, with residual reactant N being exported. In this situation, the rate of the gross supply of reactant N equals the sum of the rates of the product N generated and the residual reactant N exported. The following approximate expressions apply to the reactant N pool (d15Nreactant; Eq. 6) and the product N pool (d15Nproduct; Eq. 7) (Fig. 1): d15 Nreactant ¼ d15 Ninitial þ e
ð1 d15 Nproduct ¼ d15 Ninitial e
ðf Þ fÞ (6) (7) The special case of nearly complete consumption (f close to 0) reflects the common situation of a reactant behaving as an “intermediate”: it is introduced (here, without isotopic discrimination) from an effectively infinite N reservoir and is consumed at the same rate, with discrimination. In this case, the d15N of the reactant is equal to the d15N of the reactant supply (the “initial” in Eq. 6), plus the kinetic isotope effect of the consuming process. Many examples of this arise below. Fig. 1 Simple models of N isotope changes associated with the consumption of a substrate. The d15N of reactant and product N pools of a single unidirectional reaction as a function of the fraction of the initial reactant supply that is left unconsumed, for two different models of reactant supply and consumption, following the approximate equations given in the text. The Rayleigh model (black lines) applies when a closed pool of reactant N is consumed. The steady-state model (green lines) applies when reactant N is supplied continuously. The same isotopic parameters, an isotope effect (e) of 5% and a d15N of 5% for the initial reactant supply, are used for both the Rayleigh and steady-state models. e is approximately equal to the d15N difference between reactant N and its product (the instantaneous product in the case of the Rayleigh model). The gray dotted horizontal line indicates the initial reactant supply d15N. 266 Ocean Process Tracers | Nitrogen Isotopes in the Ocean The steady-state model and modified forms of it, such as the more spatially complex “reaction-diffusion” model, are used to quantify uptake processes where supply and uptake are simultaneous and relatively time-invariant, such as in the consumption of nitrate by denitrification in sediments. In some cases, it is not immediately clear whether the Rayleigh or steady state model is the more appropriate approximation. One example is a steady upwelling system, with a supply of nitrate by upwelling, consumption by phytoplankton assimilation, and lateral export of the remaining nitrate. Which model applies best will depend on the reference frame of the analysis (e.g., Eulerian or Lagrangian), the spatial and temporal scale of the measurements, and the N forms of greatest interest. A growing number of computer models of the ocean include the N isotopes. Two strategies exist for their implementation. First, within a time-dependent model, simple (typically first-order) expressions for N fluxes can be used. In this case, reaction rate coefficients 14k and 15k are used for the conversion rates of reactant 14N and 15N (14Nr and 15Nr) to product (t 1 and t being consecutive time steps): 14 Nr ðt Þ ¼ 14 Nr ðt 1Þ 14 k
14 Nr ðt 1Þ (8) 15 Nr ðt Þ ¼ 15 Nr ðt 1Þ 15 k
15 Nr ðt 1Þ (9) where 15 15 Nr ðt 1Þ ¼ 14 k¼ Nr ðt 14 k
1 e
10 15 3
1Þ
RAir
d Nðt (10) 1Þ
10 3
þ1 Product (14Np and 15Np) can then be calculated from the mass balance of reactant conversion:
14 Npðt Þ ¼ 14 Npðt 1Þ þ 14 Nr ðt 1Þ 14 Nr ðt Þ
15 Npðt Þ ¼ 15 Npðt 1Þ þ 15 Nr ðt 1Þ 15 Nr ðt Þ (11) (12) (13) Second, in model “boxes” where either the Rayleigh or steady-state model is adequately accurate for individual time steps, the analytical expressions for that model can be applied. Whatever the approach, especially when the N reservoir has a long residence time relative to the number of calculations, it is critical to ensure mass conservation; otherwise, isotopic error can accumulate. Processes Inputs N2 fixation is the major input of fixed N to the ocean (Fig. 2); it is carried out by “N2 fixers”, cyanobacteria and other microorganisms able to convert N2 into biomass N. Subsequent remineralization of this biomass supplies new N to the dissolved fixed N pools in the surface and subsurface ocean. Field collections of Trichodesmium colonies, the best-known genus of open ocean N2 fixer, have yielded a d15N of c. 2% to þ0.5%. Taking into account the d15N of dissolved N2 (0.6% in the surface mixed layer), this range in d15N suggests an isotope effect for N2 fixation by Trichodesmium of 2.5% or less. An average d15N of 1% has been suggested for the fixed N input to the ocean from N2 fixation, although this is almost entirely driven by d15N measurements of Trichodesmium colonies, thus not taking into account other marine N2 fixers or the exudation of DON by N2 fixers. Other inputs of fixed N to the marine environment include terrestrial runoff and atmospheric deposition, the N isotopic compositions of which are poorly constrained (Fig. 2). Dissolved and particulate N d15N measurements in pristine river systems range mostly from 0 to 5%, although this may be biased toward temperate climates, which appear to be lower in d15N than tropical systems. Biological processes along the flow path and in estuaries (in particular, denitrification) can alter the d15N of the final inputs from terrestrial runoff in complex ways. Anthropogenic inputs often increase the d15N of a system because they encourage denitrification along the path of freshwater flow. In atmospheric inputs, individual studies of conditions relatively remote from human influence have reported wide ranges in the d15N of inorganic and organic N (e.g., 6% to þ10%, 13% to 0%, and 5% to þ15%, for nitrate, ammonium, and water soluble organic N, respectively), with increasing evidence that at least some of this variability can provide insight into sources and processes. In the face of large uncertainties, a preindustrial mean d15N of 4% for terrestrial runoff and 0% for atmospheric deposition in the open ocean has been suggested. Human activities now play a central role in determining the fluxes and d15N of N in both terrestrial runoff and atmospheric deposition, probably raising the d15N of the former while clearly decreasing the d15N of the latter. Outputs Denitrification, the bacterial reduction of nitrate to N2, is the major mechanism of fixed N loss from the ocean, occurring both in the water column and in sediments where the oxygen (O2) concentration is low (<5 mmol/kg) (Fig. 2). Denitrification strongly discriminates against the heavier isotope, 15N, progressively enriching the remaining nitrate pool in 15N as nitrate consumption Ocean Process Tracers | Nitrogen Isotopes in the Ocean 267 Fig. 2 Processes affecting the nitrogen isotopes in the ocean. The inputs and outputs (dashed arrows) control the ocean’s inventory of fixed N, the majority of which is in the form of nitrate (NO3 ). Color indicates the d15N of nitrate. The low-latitude surface is colorless because of nearly complete nitrate consumption. Inputs are N2 fixation in the surface ocean, terrestrial runoff, and atmospheric deposition. Outputs indicated are sedimentary and water column denitrification. As discussed in the text, water column denitrification in low-oxygen regions at mid-depths leaves behind nitrate that is elevated in d15N (yellow area); due to limitations imposed by diffusion in porewaters, the isotopic discrimination of sedimentary denitrification is much weaker. Remineralization of newly fixed N can explain the low d15N of nitrate in the shallow subsurface, or thermocline, of the subtropical gyres (blue area), such as in the Sargasso Sea of the North Atlantic. Internal cycling is represented with solid arrows. Nitrate supplied from deep water, mid-depths and the thermocline is assimilated in the surface ocean. We draw a distinction between the high and low latitudes, represented from left to right by the Southern Ocean, the low-latitude ocean (using data from the Sargasso Sea for illustration; e.g., Fig. 7), and the coastal zone. Partial nitrate assimilation in the Southern Ocean at high latitudes increases surface nitrate d15N and is subducted into the ocean interior with deep and intermediate water mass formation (light green area), transmitting high nitrate d15N into the low-latitude, mid-depth ocean. In the low-latitude ocean, nitrate is completely consumed at the surface, such that sinking PN d15N is similar to upwelled nitrate d15N. PN is recycled in the surface ocean, being degraded to ammonium (NHþ 4 ) that is subsequently assimilated. The breakdown of DON to ammonium occurs with greater isotopic discrimination than the production of DON from PN, causing the d15N of DON to be higher than that PN. Sinking and remineralization (via degradation and nitrification, see text) returns N from the particulate pool to nitrate in the ocean interior. The isotope discrimination associated with nitrification in the ocean interior is excluded because this process generally goes to completion (see text). Question marks indicate the greatest uncertainties, due to limited data or great isotopic variation. proceeds. Looking across culture and field studies, the preponderance of data suggests an isotope effect between 15% and 25%. The isotopic discrimination during denitrification likely originates inside the denitrifier cell, with the reduction of nitrate to nitrite by the dissimilatory form of the enzyme nitrate reductase. Unconsumed, 15N-enriched nitrate effluxes from the cell back into ambient waters, allowing the enzyme-level isotope effect to be expressed in the ambient water surrounding the organism (Fig. 3). Where it occurs in low-oxygen regions of the mid-depth ocean, water column denitrification causes a clear elevation in the d15N of nitrate, and it is the main reason that global ocean nitrate d15N is higher than that of the N from N2 fixation, the dominant input. In contrast to water column denitrification, denitrification in sediments leads to little increase in the d15N of water column nitrate. The high d15N of nitrate within the pore waters of actively denitrifying sediments demonstrates that, as with denitrification in general, isotopic discrimination occurs at the scale of the organism. However, expression of the organism-scale isotope effect at the scale of sediment/water exchange is minimized by nearly complete consumption of the nitrate at the site of denitrification within sediment pore waters. This prevents 15N-enriched residual nitrate from escaping to the overlying water column, yielding an apparent isotope effect of 3% or less emanating from most sediments studied so far. The high degree of nitrate consumption in sediments is itself most often due to the slowness of exchange of pore waters with the overlying water column. With regard to the general lack of isotopic discrimination by N loss in sediments, an important exception has been identified in highly productive ocean margin environments such as the subarctic and Arctic shelves. When a rapid flux of organic N to the sediment is remineralized to ammonium, part of this ammonium is oxidized to nitrite and nitrate, much of which is subsequently denitrified. However, the remaining excess of ammonium escapes into the overlying water column. Because of isotopic fractionation during ammonium oxidation (see below), the ammonium that escapes into the water column can be elevated in d15N relative to that of the sinking flux, and the ammonium will eventually be completely converted to nitrate. When this combination of processes is integrated to yield a single apparent isotope effect, values as high as 8% have resulted. 268 Ocean Process Tracers | Nitrogen Isotopes in the Ocean Fig. 3 Cellular mechanism of isotopic fractionation associated with nitrate assimilation and denitrification. For both processes (A), the first irreversible step of nitrate reduction (NR) is the main fractionating step; a study of nitrate assimilation found that eNR is 27%. In the case of nitrate assimilation, the nitrite (NO2 ) produced by the nitrate reduction is eventually converted to tissue (particulate N, PN); for denitrification, it is eventually converted to N2. However, these following steps do not affect the isotope systematics of nitrate consumption and are thus shaded gray. The degree to which this intracellular isotope fractionation is expressed outside the cell depends on the extent of the efflux (“Out”), which allows high-d15N intracellular nitrate to be released into the extracellular environment. A small and relatively similar isotope effect is observed for both influx (“In”) and efflux. The much higher isotope effect reported for denitrification (C) than for nitrate assimilation (B) is thought to arise from the fact that, in denitrification relative to nitrate assimilation, a much larger fraction of nitrate influx is effluxed rather than reduced to nitrite and sent on for further reaction. In (B) and (C), arrow line thickness indicates relative rate. Figure concept and eNR estimate from Karsh, K. L., Granger, J., Kritee, K., and Sigman, D. M. (2012). Environmental Science & Technology 46, 5727–5735. In addition to denitrification, another mechanism of fixed N loss that occurs in sediments and the water column is anaerobic ammonium oxidation, or “anammox”, in which nitrite (typically from nitrate reduction) is used to oxidize ammonium to N2 (NO2 þ NHþ 4 ! N2 þ 2H2O). The significance of anammox in the ocean has only recently been established; the first sign of its occurrence was the lack of ammonium accumulation in the suboxic (oxygen-deficient) zones of the ocean water column. The isotopic behavior of anammox is currently being investigated. The net effect of anammox on the N isotopes will depend on the organism-scale isotope effects, the sources of nitrite and ammonium substrates for the reaction, and the degree to which these substrates are consumed. Consider, for example, the following conditions: (1) nitrate reduction by denitrifiers is the source of the nitrite, (2) nitrite reoxidation to nitrate is slow, (3) remineralization processes are the source of the ammonium, and (4) both the nitrite and ammonium are completely consumed within the suboxic zones. Under these conditions, anammox would serve solely as a loss mechanism for nitrite and ammonium, the isotope systematics of which would have no consequence, and the isotopic discrimination of N loss would simplify to that of the nitrate reduction by denitrifiers, with a minor influence on the net isotopic discrimination from the remineralization that produces the needed ammonium. However, these simplifying conditions may be violated substantially in ocean suboxic zones. For example, a side reaction of anammox produces nitrate, affecting the d15N of the total nitrate pool in a water parcel. Moreover, within the major suboxic zones, there is isotopic evidence for substantial rates of the oxidation of nitrite to nitrate, due to anammox and/or other processes. As with anammox, the recent evidence for extensive nitrite oxidation in the suboxic zones is causing a reconsideration of the isotope dynamics associated with ocean N loss. It should be noted that many of the early water-column-derived isotope effect estimates for “denitrification” regressed the nitrate þ nitrite d15N increase against the total nitrate þ nitrite deficit as calculated from phosphate, with ammonium observed not to accumulate. Accordingly, while these estimates did not explicitly consider anammox and/or nitrite oxidation, they inherently included them by integrating across all processes. Nevertheless, early estimates suffered from relying on comparison of the suboxic zone measurements with measurements from deep water below the suboxic zone. It is now clear that the suboxic samples must be compared with water samples from outside the suboxic zone that are of the same density, as only water of the same density can be interpreted as the precursor of the water in the suboxic zone. Accordingly, readers seeking water column denitrification isotope effects should tend to focus on recent studies. Internal Cycling The fluxes associated with internal cycling are neither sources nor sinks of oceanic fixed N, but they affect the distributions of N species and isotopes in the ocean. Nitrogen assimilation In the surface ocean, phytoplankton assimilate fixed N (nitrate and ammonium, as well as nitrite and various organic N compounds). Culture studies indicate that different forms of fixed N are assimilated with distinct isotope effects (Fig. 4), Ocean Process Tracers | Nitrogen Isotopes in the Ocean 269 Fig. 4 Isotope effects (e values) of processes central to the internal cycling of N in the ocean. The isotope effects shown here are based on laboratory and field studies. Dashed arrows represent assimilation of dissolved species into particulate matter, and solid arrows represent remineralization. Complete consumption of the ammonium pool by assimilation in the surface ocean or by nitrification in the ocean interior causes the relatively high isotope effects associated with these processes to have little effect on N isotope dynamics. However, in regions where ammonium assimilation and nitrification co-occur, their isotope effects will impact the d15N of their respective products, PN and nitrate. In nitrification, ammonia (NH3), rather than the protonated form ammonium (NHþ 4 ), is oxidized. However, ammonium is the dominant species in seawater, and there is isotope discrimination in the ammonium–ammonia interconversion. Thus, the isotope effects for “ammonia oxidation” given here and in the text refer specifically to consumption of ammonium. Relative to the breakdown of PN to DON, the breakdown of DON to ammonium appears to occur with a larger isotope fractionation (roughly 5%), which acts to raise the d15N of DON relative to PN. It may be that most PN breakdown to ammonium occurs through DON. The isotope effect estimates in this figure derive mostly from Casciotti, K. L., Sigman, D. M., and Ward, B. B. (2003). Geomicrobiology Journal 20, 335–353; Casciotti, K. L. (2009). Geochimica et Cosmochimica Acta 73, 2061–2076; Knapp, A. N., Casciotti, K. L., and Prokopenko, M. G. (2018). Global Biogeochemical Cycles 32, 769–783; Möbius, J. (2013). Geochimica et Cosmochimica Acta 105, 422–432; Vo, J., Inwood, W., Hayes, J. M., and Kustu, S. D. (2013). PNAS 110(21), 8696–8701; Waser, N. A. D., Harrison, P. J., Nielsen, B., Calvert, S. E., and Turpin, D. H. (1998). Limnology and Oceanography 43(2), 215–224. although these isotope effects may vary with physiological conditions. For all studied forms, phytoplankton preferentially consume 14 N relative to 15N. Nitrate is the subsurface-water source of fixed N for phytoplankton growth. The degree of its consumption varies across the surface ocean, with essentially complete assimilation by phytoplankton in the stably stratified tropics and the subtropical gyres but only partial consumption in polar and upwelling regions (Fig. 2). Certain polar surface regions responsible for injecting water into the mid-depths or deep ocean (in particular, the Southern Ocean) are characterized by incomplete nitrate consumption. The middepth nitrate, part of which derives from the Southern Ocean surface, then fuels productivity throughout the low-latitude ocean when it is upwelled or mixed to the surface. Because of these and other processes, the discrimination against 15N during nitrate assimilation has a major impact on the isotopic distributions of all N forms in the ocean. Field-based estimates of the isotope effect of nitrate assimilation range from 4% to 10%, with most estimates between 4% and 7%. Culture-based estimates are more variable, in some cases for identifiable reasons. Physiological studies suggest that isotopic fractionation associated with nitrate assimilation is imparted by the intracellular assimilatory nitrate reductase enzyme, which has an intrinsic isotope effect of c. 27%. As with denitrification, the enzyme-level isotope effect is expressed by efflux of unconsumed nitrate out of the cell (Fig. 3). The lower range of the isotope effect associated with phytoplankton nitrate assimilation than with denitrification suggests proportionally less nitrate efflux in the former (Fig. 3B and C). This difference may be related to the fact that N limitation of phytoplankton is common, which may have caused nitrate assimilation to evolve to minimize nitrate efflux. The degree of efflux and isotope effect of nitrate assimilation may vary with growth conditions; laboratory studies and some field incubations indicate a higher isotope effect under light-limited growth than under growth limited by iron or temperature. Other forms of fixed N assimilated by phytoplankton (ammonium, nitrite, and urea; Fig. 4) are produced and nearly completely consumed in the open ocean surface mixed layer, with a balance between production and consumption on time scales of days or less. Therefore, the isotopic discrimination associated with the assimilation of these N forms has less impact on N isotope distributions than in the case of nitrate. Culture and field studies suggest an isotope effect for ammonium assimilation of up to c. 20% under high ammonium concentrations, but decreasing to c. 0%–4% as ammonium concentration decreases. The relationship with ammonium concentration is thought to be related mostly to the mode of transport across the cell membrane: active NHþ 4 transport versus passive (diffusive) NH3 transport, the latter incorporating the high (c. 19%) equilibrium isotope effect between NHþ 4 and NH3. Minimal isotope effects (<1%) have been reported for assimilation of nitrite and urea, possibly due to lack of efflux of these species once they have been imported into the intracellular environment. 270 Ocean Process Tracers | Nitrogen Isotopes in the Ocean Remineralization The return of organic N to nitrate occurs in two steps: (1) the degradation of organic N to ammonium and (2) the bacterial oxidation of ammonium to nitrate, or “nitrification” (Fig. 4). Nitrification itself occurs in two steps, the oxidation of ammonium to nitrite and the oxidation of nitrite to nitrate, mediated by distinct groups of microorganisms. Isotopic discrimination may occur at all steps involved in remineralization. Field studies generally suggest that both bacteria and zooplankton preferentially degrade lowd15N organic N, yielding residual organic matter relatively high in d15N. The wide spectrum of reactions involved in organic N degradation and the heterogeneous nature of organic matter (comprised of compounds with distinct d15N that degrade at various rates by different pathways) make quantifying the isotope effect associated with degradation difficult. A few laboratory studies have quantified the isotope effects of individual processes such as thermal peptide bond cleavage, bacterial amino acid uptake and transamination, deamination, and zooplankton ammonium release. Laboratory studies attempting to mimic degradation as a whole and field studies investigating the relationship between the degree of remineralization and residual organic matter d15N suggest a net isotope effect of 3%. Culture studies indicate a large isotope effect for the conversion of ammonium to nitrite, the first step in nitrification. Estimates of the isotope effect for marine ammonia-oxidizing bacteria and archaea range from 13% to 41%. The isotope effect of nitrification estimated from ammonium concentration and d15N measurements in the Chesapeake Bay is 12%–16%, in the lower range given by the culture studies. In culture studies of nitrite-oxidizing bacteria, the conversion of nitrite to nitrate, the second step in nitrification, is observed to occur with an inverse kinetic isotope effect (c. 13%), preferentially oxidizing 15N-nitrite to nitrate. Given the isotopic discrimination associated with each of the major steps that converts organic N to nitrate as well as the fact that essentially all organic N produced in the ocean is eventually nitrified, it would be intuitive to expect this conversion to cause large scale isotopic variations in ocean nitrate. However, its overall isotopic footprint is far less than this. The reason is that, on adequately large scales of time and space, the conversion of organic N to nitrate approaches completion, such that the d15N of nitrate is nearly equal to the d15N of the organic N undergoing remineralization. Beyond this, some opportunities for an isotopic impact of remineralization appear to be missed. For example, the evidence for isotopic fractionation during organic N degradation might lead to the expectation that sinking PN would preferentially release low d15N N soon after leaving the surface, lowering the nitrate d15N of the upper water column relative to the deep ocean. However, as discussed below, sinking PN d15N does not appear to increase with depth, arguing against an isotopic impact associated with N loss from sinking PN. As described below, the increase with depth in the d15N of suspended PN (i.e., PN that can be filtered from the water and is mostly too small and/or buoyant to sink) is consistent with little isotopic discrimination during the transformation of sinking to suspended PN, followed by significant isotopic discrimination but complete conversion during the degradation of suspended PN to ammonium. In this case, the d15N of newly nitrified nitrate will approximate the d15N of the N sinking out of the surface ocean. Nitrogen Reservoirs Dissolved Nitrogen Nitrate Nitrate accounts for most of the fixed N in the ocean. The d15N of nitrate in the deep ocean (below 1500 m) is typically between 4.5% and 5.5%. Regionally, the d15N of nitrate varies between roughly 2% and 25% due to the effects of N2 fixation, nitrate assimilation, and denitrification (Fig. 5A), but most measurements fall within a much narrower range (Fig. 6). Nitrate d15N significantly lower than deep-ocean nitrate has been observed in the upper thermocline of the low-latitude oligotrophic ocean (Figs. 2, 6C and 7). This 15N depletion is probably mostly due to the oxidation of newly fixed N, which, as described above, has a d15N of c. 1%. Values higher than 5% mostly result from discrimination associated with nitrate assimilation by phytoplankton at the ocean surface (Fig. 6B and D) or denitrification in oxygen-deficient zones of the ocean interior (Fig. 6D). Below 2 km depth, among the different ocean basins, nitrate d15N is relatively constant at c. 5%, despite large inter-basin differences in nitrate concentration. The minimal degree of isotopic variation in the nitrate of the deep ocean is due to the fact that, in most surface waters, the nitrate supply from below is almost completely consumed by phytoplankton, such that the organic N exported from the surface ocean converges on the d15N of the nitrate supply. Because the sinking flux d15N is close to that of the nitrate supplied from the ocean interior, remineralization of the sinking flux in the ocean interior does not alter greatly the d15N of deep nitrate (Fig. 5A). In this respect, the oceanic cycling of the N isotopes differs markedly from that of the carbon isotopes. There are regions of the ocean where nitrate is incompletely consumed in surface waters, such as the high-latitude, nutrient-rich regions of the Southern Ocean and the subarctic Pacific, and the low-latitude upwelling regions of the California Current and the Equatorial Pacific. In the surface waters of these regions, nitrate d15N is elevated (Fig. 6B and D), with a strong correlation between the degree of nitrate consumption by phytoplankton and the d15N of the nitrate remaining in the water. This surface signal of nitrate consumption is central to the paleoceanographic application of the N isotopes to reconstruct nutrient conditions in the surface ocean at times in the past. The polar and subpolar surface ocean provides the water that fills the mid-depth ocean (the “intermediate” waters at 1000 m depth and the “mode” and thermocline waters above). Partial consumption by assimilation elevates the d15N of nitrate in the polar and subpolar surface ocean and thus causes modest but important d15N elevation in most newly formed mid-depth waters (Figs. 2 and 6B). Due to this effect, the d15N of organic N produced in and exported from much of the low-latitude surface ocean is 1%–3% Ocean Process Tracers | Nitrogen Isotopes in the Ocean 271 Fig. 5 The effect of different marine N cycle processes on nitrate d15N and concentration ([NO3 ]) (A), nitrate d18O and concentration (B) and nitrate d18O and d15N (C), assuming an initial nitrate concentration, d15N and d18O of 30 mmol/kg, 5% and 2%, respectively. The trajectories are for typical estimates of the isotope effects, and they depend on the initial nitrate d15N and d18O as well as the relative amplitude of the changes in nitrate concentration (30% for each process). A dashed arrow denotes a process that adds or removes fixed N from the ocean, while a solid arrow denotes a component of the internal cycling of oceanic fixed N. The effects of these two types of processes can be distinguished in many cases by their effect on the concentration ratio of nitrate to phosphate in seawater. The actual impact of the different processes on the N isotopes varies with environment. For instance, if phytoplankton completely consume the available nitrate in a given environment, the isotope effect of nitrate assimilation plays no major role in the d15N of the various N pools and fluxes. Similarly, the lack of a large isotope effect for sedimentary denitrification is due to the fact that nitrate consumption by this process can approach completion within sedimentary pore waters. The d15N of nitrate regenerated in the ocean interior (the solid green arrows to the right) varies greatly (green shading) due to its dependence on the d15N of the organic N being remineralized. The d15N of the organic N sinking into the ocean interior depends on the d15N of nitrate supplied to the surface, the effects of partial nitrate consumption, and the relative contribution of N2 fixation to the organic N (dashed green arrow). In contrast, the d18O of regenerated (“newly nitrified”) nitrate is insensitive to whether the ammonium fueling the nitrification derives from organic N produced by partial nitrate assimilation, complete nitrate assimilation, or N2 fixation. The main control identified so far is the d18O of ambient seawater. Figure courtesy of Dario Marconi. Fig. 6 Three nitrate d15N depth sections through the ocean, each dominated by the isotopic signal of an important process in the ocean N cycle. (A) The section locations, overlain on surface nitrate concentration (in mmol/kg). Colors indicate nitrate d15N, with the same scales in B and C but an expanded scale in D; white isolines indicate nitrate concentration. (B) The roughly meridional GOSHIP IO8S section through the Southern Ocean shows an increase in d15N toward the surface due to partial consumption of the nitrate by phytoplankton assimilation. At 40 –50 S, Southern Ocean surface water is subducted into the upper 1000 m, ventilating the lower latitude upper ocean with high-d15N nitrate. (C) The northwest/southeast GEOTRACES GA03 section through the North Atlantic shows low-d15N nitrate in the shallow thermocline, reflecting N2 fixation in the equatorial and tropical Atlantic. (D) Zonal section GEOTRACE GP16 in the eastern South Pacific shows a large increase in residual nitrate d15N at mid-depth (125–300 m) in the easternmost stations due to partial nitrate consumption by water column denitrification. d15N elevation at the surface in the western stations due to the partial assimilation of upwelled nitrate along the equator is also evident. Data from Marconi, D., Weigand, M. A., Rafter, P. A., McIlvin, M. R., Forbes, M., Casciotti, K. L., and Sigman, D. M. (2015). Marine Chemistry 177, 143–156; Peters, B. D., Lam, P. J., and Casciotti, K. L. (2018). Marine Chemistry 201, 137–150; Fripiat, F., Martínez-García, A., Fawcett, S. E., Kemeny, P. C., Studer, A. S., Smart, S. M., Rubach, F., Oleynik, S., Sigman, D. M., and Haug, G. H. (2018). Submitted to Geochimica et Cosmochimica Acta. Figure Courtesy of Dario Marconi. 272 Ocean Process Tracers | Nitrogen Isotopes in the Ocean higher than it would otherwise be. In contrast, the downward transport of partially consumed surface nitrate by deep water formation around Antarctica does not appear to cause significant regional increases in the d15N of nitrate in the deep ocean, in part because the sinking N produced by the partial nitrate assimilation, with its low d15N of 0%–3%, is ultimately remineralized back to nitrate in polar deep waters. Indeed, if partial nitrate assimilation elevates the d15N of mid-depth nitrate on a global basis, then by mass balance, the low-d15N sinking N produced during this partial nitrate assimilation must lower the d15N of global deep (subintermediate water) nitrate, even if by only a small amount. The regional manifestation of this is found in the deep Antarctic Ocean, which has the lowest nitrate d15N measured in the global deep ocean (reaching down to 4.5%). The d15N organic N produced by N2 fixation is c. –1%, which is only slightly depleted in 15N relative to dissolved N2 (0.6%) but significantly depleted in comparison to mean ocean nitrate (c. 5%). Sinking PN and subsequent oxidation to nitrate transmit this low-d15N signal to the subsurface nitrate pool. The isotopic imprint of N2 fixation on nitrate is strongest in the shallow subsurface (i.e., thermocline) of the subtropical gyres (Fig. 2), probably mostly due to the low background concentration of nitrate at these depths compared to the deeper ocean. In the North Atlantic subtropical gyre, shallow subsurface nitrate d15N can reach values below 2% (Figs. 6C and 7). It should be mentioned that this 15N depletion may have a contribution from isotopic fractionation during remineralization in the euphotic zone and the shallow subsurface, but this process remains to be addressed quantitatively. Because water column denitrification occurs in the subsurface and consumes only a fraction of the nitrate available, its isotope effect is efficiently expressed in the d15N of ocean nitrate at both the regional and global scales. In denitrifying regions of the water Fig. 7 Compilation of N isotope data from the Sargasso Sea near the Bermuda Atlantic Time-series Study Station. (Lower panels) Compilation of vertical profiles for the concentrations of nitrate þ nitrite and the d15N of nitrate. Gray symbols represent nitrate with a concentration lower than 1.2 mmol/kg (note the scale break at 1 mmol/kg), as these nitrate samples are strongly affected by nitrate assimilation, with preferential 14N consumption causing their elevation in d15N. Average concentration and d15N profiles are also shown for total (dissolved þ particulate) organic N (TON); DON can be calculated by difference from the much smaller pool of suspended PN. (Upper panels) Compilation of particulate N concentration and d15N for sinking PN, suspended PN, zooplankton, and flow cytometrically sorted or manual picked components of the microbial community. The boxes and whiskers cover the 25th to 75th and the 10th to 90th percentile, and the line within the box indicates the median of the observations. No concentration is reported for the analyzed Trichodesmium colonies or zooplankton; sinking PN is reported as a rate and is not given here. Trichodesmium data predominantly from Carpenter, E. J., Harvey, H. R., Fry, B., and Capone, D. G. (1997). Deep-Sea Research I 44(1), 27–38. Sinking, PN data from Altabet, M. A. (1988), Deep-Sea Research 35(4), 535–554. Zooplankton and remaining sinking PN data from Smart, S. M., Ren, H., Fawcett, S. E., Schiebel, R., Conte, M., Rafter, P. A., Ellis, K.K., Weigand, M. A., Oleynik, S., Haug, G. H. and Sigman, D. M. (2018). Geochimica et Cosmochimica Acta 235, 463–482. TON data from Knapp, A. N., Sigman, D. M., and Lipschultz, F. (2005), Global Biogeochemical Cycles 19, GB1018, doi:10.1029/2004GB002320. Data for nitrate and the remaining particulate N forms from Fawcett, S. E., Lomas, M. W., Casey, J. R., Ward, B. B., and Sigman, D. M. (2011). Nature Geoscience 4, 717–722, and Fawcett, S. E., Ward, B. B., Lomas, M. W., and Sigman, D. M. (2015). Deep-Sea Research I 103, 64–72. Ocean Process Tracers | Nitrogen Isotopes in the Ocean 273 column, the d15N of nitrate commonly reaches above 15% (Fig. 6D). The subsurface d15N maximum occurs in the core of the suboxic (O2-deficient) zone and is correlated with the degree of nitrate consumption by water column denitrification. Denitrification, both in the water column and sediments, exerts control on the d15N of mean deep ocean nitrate. When the ocean N budget is at steady state, the d15N of the fixed N removed (through water column and sedimentary denitrification) will equal the d15N of the fixed N added (c. 1%, approximating N2 fixation as the sole source) (Fig. 8). If denitrification with an isotope effect of 20% were occurring homogenously in the ocean water column and responsible for all fixed N loss, the d15N of mean oceanic nitrate would be 19% to achieve a d15N of 1% for N loss. That the modern mean oceanic nitrate d15N is only 5% reflects at least two factors: (1) the importance of sedimentary denitrification, which appears to express a minimal isotope effect and (2) the localized nature of water column denitrification. With regard to the second, denitrification can consume a significant fraction of the ambient nitrate in the ocean’s suboxic zones. This water with high nitrate d15N but lowered nitrate concentration is mixed with the unfractionated and nitrate-replete pool in the surrounding ocean interior. Because the latter is somewhat higher in nitrate concentration than the former, it will give more weight to the resulting mixing product, reducing the expression of the organismlevel isotope effect of water column denitrification and thus lowering the mean d15N of nitrate required to achieve an isotope balance between inputs and outputs. With these considerations, one study estimates that water column denitrification is responsible for 30% of fixed N loss from the modern ocean, with sedimentary denitrification responsible for the remainder. However, the isotope-based budget for marine fixed N remains highly uncertain. One limitation of the N isotopes as a tool to investigate N cycling in the ocean is their inability to separate co-occurring processes with competing N isotopic signatures, such as (1) denitrification and N2 fixation or (2) nitrate assimilation and regeneration/ nitrification. Coupled analysis of N and O isotopes in nitrate can help to disentangle such otherwise overprinting processes (Fig. 5). Culture studies have demonstrated that the two most important nitrate-consuming processes, nitrate assimilation and denitrification, fractionate the N and O in nitrate with a ratio close to 1:1. In contrast, the remineralization of organic N to nitrate has different effects on the nitrate N and O isotopes. The d15N of the newly produced nitrate is set by the organic matter being remineralized, taking into account the discussion above regarding isotopic discrimination during remineralization. This d15N is, in turn, sensitive to the d15N of the nitrate supplied from the subsurface, the degree of nitrate consumption in the euphotic zone, and N2 fixation. In contrast, the d18O of nitrate appears to track (with a positive offset of c. 1.1%) the d18O of the water in which the remineralization is occurring; given the overall homogeneity of water d18O in the ocean, the d18O of newly produced nitrate is expected to vary little. Due to these differences, deviations in the d18O and d15N in nitrate from 1:1 variation can provide information about the d15N of nitrate being added by remineralization and thus of the N exported from the euphotic zone. Fig. 8 The sensitivity of global mean ocean nitrate d15N to the relative importance of water column versus sedimentary denitrification, illustrated in terms of the requirement (at steady state) for the isotopic impacts of N2 fixation and denitrification to cancel one another. It is assumed here that, for all three cases shown, the ocean N budget is balanced by equal fluxes of N2 fixation and denitrification (i.e., their summed horizontal components are equal). For a stable mean ocean nitrate d15N to be achieved, the vertical components (i.e., their net isotopic impacts) must also be equal. This requirement determines the d15N of mean ocean nitrate (the vertical position of the black circles) for the different proportions of water column and sedimentary denitrification considered. N2 fixation adds nitrate with a d15N of c. 1% (blue arrows). Sedimentary denitrification removes nitrate with a d15N similar to that of mean ocean nitrate (brown arrows). In one hypothetical end-member case (dotted arrows), it is assumed that ocean N loss is entirely by sedimentary denitrification. In this case, the isotopic impacts of input and output are balanced when the d15N of global ocean nitrate is the same as that of N2 fixation. At the opposite end-member (dashed arrows), water column denitrification is responsible for all ocean N loss. Water column denitrification removes nitrate with its characteristically large isotope effect (red arrows), such that the isotopic impacts of N2 fixation and denitrification are balanced when nitrate d15N is high (19% given the assumptions made here). The observed ocean nitrate d15N of c. 5% (solid arrows) implies a partitioning between water column and sedimentary denitrification in which sedimentary denitrification is the greater part of the total N loss. The lengths of the vectors are for intercomparison only; the absolute lengths are arbitrary. Figure courtesy of Dario Marconi. 274 Ocean Process Tracers | Nitrogen Isotopes in the Ocean Nitrite Nitrite is an intermediate in oxidative and reductive processes such as nitrification (Fig. 4) and denitrification. It also serves as a substrate for anammox and can be assimilated by phytoplankton. Nitrite is only observed to accumulate to significant levels when these processes become uncoupled, such as (1) in oxygen-deficient zones (with nitrite concentrations up to 10 mmol/kg, representing as much as 25% of the combined NO3 þNO2 pool), (2) at the base of the euphotic zone in most regions (with 0.05–1.5 mmol/ kg), and (3) in the surface waters of some polar regions (with c. 0.25 mmol/kg). The d15N and d18O of nitrite are starting to be systematically measured in the ocean, a currently active area of research. Limiting discussion here to the d15N of nitrite, it is expected to reflect the balance of isotopic discrimination during nitrite production and consumption processes (Fig. 4). The d15N of nitrite is low ( 100% to 0%), 5%–105% lower than the nitrate in the same water, and varies strongly with depth. In the oxygen-deficient zones, nitrite d15N is lower than expected from denitrification alone, given the known isotope effects for nitrate and nitrite reduction. This low d15N likely points to other processes acting on the nitrite pool, with nitrite oxidation and nitrate/nitrite isotope equilibration both being investigated. The same processes may also apply in certain oxic environments, such as surface waters of the Southern Ocean, where nitrite d15N is also lower than expected. In these environments, nitrite significantly impacts the measured d15N of the combined NO3 þNO2 pool. Until recently, the two species nitrate and nitrite have typically been combined in isotopic analysis, such that the presence of low-d15N nitrite may have masked some of the 15N enrichment of nitrate in oxygen-deficient zones and sunlit surface waters. Moreover, nitrite remains reactive under most methods of seawater sample preservation, with potentially complex isotopic repercussions; its effects on nitrate isotopic analyses both within suboxic zones and in the oxic upper ocean are currently being studied. The methods for nitrite removal (with or without nitrite isotopic analysis) can be applied to investigate and/or remove the impact of nitrite on the measured d15N of nitrate. Ammonium The d15N of ammonium reflects the production of ammonium by the degradation of organic N and its consumption by nitrification, ammonium assimilation (Fig. 4), and anammox. Analytical constraints have limited isotopic studies of ammonium to environments with ammonium concentrations greater than 1 mmol/kg, precluding studies in the open ocean. In estuarine systems, where ammonium can be abundant, its d15N is often high (commonly higher than 10%), and it increases as the ammonium concentration decreases along transects from riverine to marine waters due to isotopic discrimination associated with ammonium consumption by nitrification and/or assimilation. In the open ocean surface mixed layer, it is generally assumed that ammonium generated by remineralization is quickly and entirely assimilated by plankton, in which case the isotope effect associated with its consumption would not play an important role in N isotope dynamics of the open ocean surface. In the open ocean interior, below the depth of assimilation by phytoplankton, essentially all ammonium generated from particles is oxidized to nitrite and then nitrate before it can be transported into or out of a given region. Thus, nitrification should be of limited importance for the isotope dynamics of both particulate and dissolved N once the former has sunk out of the upper ocean. However, in at least some regions of the upper ocean, ammonium assimilation and oxidation are likely to co-occur. If the isotope effect of ammonium oxidation is greater than that of ammonium assimilation, low-d15N N will preferentially be routed to the nitrate pool by oxidation, and high-d15N N will be routed back to the PN pool by assimilation. If the isotope effect of oxidation is less than that of assimilation, the opposite will occur. Thus, with better constraints on the isotope effects of ammonium-consuming processes, the isotopes of upper ocean N pools promise to provide an integrative constraint on the relative rate of nitrification in the upper ocean, especially when paired with the O isotopes of nitrate (see above). Dissolved organic nitrogen DON concentrations are significant in the open ocean, typically 4 mmol/kg in surface waters, decreasing to c. 2 mmol/kg in deep water. Fluxes associated with the DON pool are among the least understood components of the marine N cycle. Studies to date of bulk DON have focused on subtropical surface waters, in which DON is by far the dominant N pool (Fig. 7). In the low-latitude upper ocean, the d15N of DON follows the d15N of the nitrate supply from the shallow subsurface, being higher than it by c. 1%. The d15N of upper ocean DON is thus also similar to (1% higher than) that of sinking N, while it is c. 4% higher than suspended PN (see below; Fig. 7). The breakdown of DON to ammonium would be expected to involve isotopic fractionation, and the existing field data support this, suggesting an isotope effect of 5% (Fig. 2). In contrast, the first step in the production of DON from PN likely does not involve the N itself (e.g., physical disintegration, solubilization, breakdown at carbon–carbon bonds), in which case no N isotopic fractionation should occur. Moreover, isotopic fractionation associated with N-involving bond breakage that does occur to produce DON would be diluted across all the N atoms in the newly produced DON molecule. Thus, no significant N isotopic fractionation is expected during DON production. This view of fractionation during the degradation but not the production of DON can explain the d15N elevation of total organic N (TON) relative to suspended PN in the low-latitude surface ocean (Fig. 7; lower right and upper right panels, respectively). In a sense, this explanation is analogous to that given for the d15N elevation of zooplankton relative to their food source: The zooplankton feed without a strong isotopic discrimination but metabolize low-d15N N, which is then available as ammonium to the phytoplankton. Minimal gradients in the concentration and d15N of DON in the upper ocean hinder reconstruction of the fluxes of DON and the 15 d N of those fluxes. Progress on DON d15N dynamics would be aided by a method to remove nitrate from samples without Ocean Process Tracers | Nitrogen Isotopes in the Ocean 275 compromising the DON pool, which would make subsurface waters and high-nitrate surface waters more accessible to study. Ongoing work on separable fractions of the DON pool (e.g., the high-molecular-weight fraction and its components) is also promising. Particulate Nitrogen Suspended particles A typical profile of suspended particles has its lowest d15N in the surface layer, increasing below the euphotic zone (Fig. 9). In part, the d15N of suspended particles reflects the d15N of nitrate supplied to and consumed by phytoplankton, taking into account the degree of nitrate consumption. However, it is also affected by subsequent N cycling within the euphotic zone, as well as by N2 fixation in some environments. For example, in the low-latitude, low-nutrient surface ocean (e.g., the Sargasso Sea; Figs. 2 and 7), the d15N of suspended PN in the surface layer is typically 2%–4% lower than expected solely from assimilation of the nitrate supply from below. Traditionally, this low d15N has had two competing explanations: N2 fixation and N recycling. As described above, N2 fixation adds fixed N with a d15N of c. 1% to surface waters. In most cases, however, calculated rates of N2 fixation are too low to explain the low d15N of bulk suspended PN in the euphotic zone. The isotopic effect of N recycling originates from heterotrophic processes. Zooplankton release ammonium with a low d15N, making their tissues and solid waste higher in d15N than their food source. The low d15N ammonium is consumed by phytoplankton and thus retained in the surface ocean N pool, while the 15Nenriched PN is preferentially exported as sinking particles, leading to a lower d15N of surface PN. This mechanism explains why a lower-than-expected d15N is commonly reported for suspended PN in the surface ocean, including in cold environments where significant N2 fixation is highly unlikely. Recent measurements of the d15N of flow cytometry-sorted microbial groups point to an additional mechanism for lowering the 15 d N of suspended PN in the surface ocean (Fig. 7). The cyanobacterial picoplankton taxa Synechococcus and Prochlorococcus, neither of which fix N2, have a uniformly low d15N, consistent with their dominant assimilation of low-d15N ammonium. In contrast, the d15N of small eukaryotic phytoplankton is higher and thus closer to the d15N of the nitrate supply to the euphotic zone, implying that they are responsible for most of the nitrate assimilation in the euphotic zone. Relative to the cyanobacteria, the small eukaryotes are more amenable to incorporation in sinking PN, either by packaging in fecal pellets or incorporation in sinking aggregates. Disproportionate representation of small eukaryotes in sinking PN would tend to raise its d15N, leaving lower-d15N N in the surface ocean. The d15N of suspended particles in the subsurface is typically 6% higher than suspended particles in the surface ocean and 3% higher than the sinking flux (Fig. 9). The d15N of deep particles is consistent with the inferences that deep particles are the breakdown product of material exported from the surface and that bacteria preferentially remineralize low-d15N PN. Isotopic analysis of zooplankton and organisms at higher trophic levels can provide insights into the marine N cycle. The “trophic effect”, an observed c. 3% increase in d15N per trophic level that presumably results from isotopic discrimination during metabolism of N-bearing organic matter, is used widely in food-web studies. N isotopic analysis of specific amino acids within d15N (‰) −1 0 0 1 2 3 4 5 6 7 8 9 500 Depth (m) 1000 1500 2000 2500 3000 3500 Sediment trap Suspended particles Fig. 9 Nitrogen isotopic values of suspended particulate matter and sinking particles (as collected by sediment traps) in the subtropical North Atlantic Ocean near BATS (31 500 N, 64 100 W; see Fig. 7). The profiles of suspended PN show the representative depth gradient in d15N, with lower d15N in the surface ocean than at depth. The d15N of the sinking flux shows a decrease with depth, which has also been observed in other regions. Reprinted from Altabet, M. A., Deuser, W. G., Honjo, S., and Stienen, C. (1991), Nature 354, 136–139. 276 Ocean Process Tracers | Nitrogen Isotopes in the Ocean organisms and particulate organic matter promises new insights as some amino acids increase in d15N with trophic level while others preserve the d15N of the food source. For example, amino acid-specific d15N measurements in deep suspended particles support the view that deep suspended PN derives from sinking PN but has undergone bacterial degradation with isotopic discrimination. Sinking particles Because vertical sinking is an important mode of N export from the surface ocean, the d15N of the sinking flux is one of the most valuable N isotopic constraints on modern ocean processes. Combined with other isotopic data, sinking flux d15N data can provide information on the routes and mechanisms of nitrate supply and can be used to constrain other sources of N to the surface. The sinking flux also transfers the isotopic signal from the surface ocean to the seafloor, providing the link through which the sediment column records the history of surface ocean processes. Sinking particles collected in depth arrays of sediment traps often show a modest decrease in d15N with depth (Fig. 9). Given the evidence for isotopic discrimination during organic N breakdown (e.g., as apparent in deep suspended PN), the absence of an increase in sinking particle d15N with depth calls for explanation. One interpretation of the lack of d15N rise is that sinking PN flux declines with depth mostly because of the disaggregation of large sinking particles into smaller suspended particles. While this disaggregation likely involves microbial degradation of the particles, it would appear that this degradation mostly does not involve the breakage of chemical bonds that include N, so that sinking particles are lost without N isotopic fractionation. The suspended particles released into the water by such disaggregation will eventually be degraded to ammonium and then nitrified to nitrate in the water in which the disaggregation occurs. If so, then there is no net isotope fractionation in the conversion of sinking PN to deep ocean nitrate. This explanation does not address, however, why some data indicate a decline in sinking PN d15N with depth (as in Fig. 9), for which there is as yet no compelling explanation. Nitrogen in sediments and the sedimentary record The controls on sedimentary N d15N are of interest primarily with regard to efforts to reconstruct past N cycle changes. There is generally a good correlation between the d15N of surface sediments and that of sinking PN from the overlying water column. In regions of the ocean where a relatively large fraction of the organic rain is preserved in the sediment column, as occurs along continental margins, this correlation is generally excellent, and there is no clear d15N difference (i.e., no offset) between sinking and sedimentary N. In open ocean sediments where only a very small fraction of N is preserved, while spatial patterns in the d15N of sediment core tops generally reflect those in sinking PN, a significant 15N enrichment (of 2%–5%) is observed in the sediment N relative to sinking PN. Upon burial, reactions in the shallow sediment column known collectively as “diagenesis” can cause a clear increase in the d15N of PN as it is incorporated into the sediment mixed layer, and this process is likely one of the main drivers of variability in the d15N relationship between sinking and sedimentary N. The input of N from land, the continental shelves or other non-local sources can also be significant in some deep ocean settings, for example, within marginal basins or on continental slopes. To address concerns regarding alteration of both sinking and sedimentary bulk d15N, studies of the past N cycle are increasingly focusing on isolating specific N components, the d15N of which is insensitive to diagenesis. These “N paleoproxies” include N protected within the mineral matrix of microfossils and fossils as well as forms of N that do not change in d15N as they are degraded, such as chlorophyll-derived compounds. While the total N sinking out of the euphotic zone is constrained by mass balance to approximate the d15N of the N supply to the euphotic zone, this does not apply to any sub-fraction of the sinking N, including the N associated with these novel paleoproxies. For example, the d15N of foraminifera shell-bound organic matter may vary as a function of the organism’s trophic level and whether they have algal symbionts. For such proxies, modern ocean groundtruthing is particularly important. The number, type, areal coverage, and time range of N isotope records are increasing rapidly. The processes and parameters reflected in these reconstructions include (1) mean ocean nitrate d15N, which is described above as sensitive to the relative importance of fixed N loss from the water column versus the sediments, (2) regional subsurface nitrate 15N depletion or enrichment relative to the global ocean owing to N2 fixation or water column denitrification, (3) regional isotope dynamics associated with partial nitrate assimilation in surface waters, and (4) possible direct contributions of newly fixed N (by in situ N2 fixation or atmospheric N deposition) to the euphotic zone. Paleoceanographers have focused on bulk sediment or paleoproxy d15N records underlying environments where a single process or parameter is thought to dominate d15N changes. For example, in the tropical North Atlantic, N2 fixation lowers the d15N of nitrate in the thermocline relative the deep ocean value (Figs. 2, 6B, and 7), such that foraminiferal d15N changes have been interpreted in terms of the waxing and waning of regional N2 fixation. Foraminifera-bound N isotope records from these regions yield higher d15N during ice ages, suggesting a weaker lowering of nitrate d15N in shallow thermocline nitrate due to a reduction in N2 fixation rates during ice ages (Fig. 10, blue records). In denitrifying regions, bulk sediment d15N changes have been taken to largely reflect changes in regional 15N enrichment due to water column denitrification. These records show lower d15N during the ice ages, suggesting reduced rates of denitrification (Fig. 10, red record). Together with the apparent reduction in N2 fixation, this would suggest a more sluggish input/output budget for oceanic fixed N during the ice ages. In high-nutrient regions, d15N records have been interpreted in terms of the degree of nitrate consumption by phytoplankton assimilation, pointing to changes in the proportion of the gross nitrate supply to surface waters that is assimilated by phytoplankton and exported from the surface as sinking PN. Diatom-bound and foraminifera-bound d15N records from the Southern Ocean indicate higher d15N during the ice ages (Fig. 10, green records). This higher d15N argues for more complete nutrient consumption, findings that are potentially important for explaining the reduction of atmospheric CO2 levels during ice ages. Fig. 10 d15N records from ocean sediments spanning the last 160 thousand years, which encompass recent cold glacial intervals (shaded) and warm interglacials (unshaded). Blue: Planktonic foraminifera-bound d15N records from the South China Sea (B) and the Caribbean Sea (C) indicate changes in the d15N of nitrate of the thermocline, suggesting lower N2 fixation rates during glacials. Red: Bulk sedimentary d15N record from the eastern tropical North Pacific (D), suggesting lower water column denitrification rates during glacials. Green: Foraminifera- (E) and diatom-bound (F) d15N records from the Southern Ocean, with high d15N suggesting more complete assimilation of the nitrate supply during glacials. The glacial and interglacial intervals were delimited on the basis of the record of glacial ice volume on land, which is plotted as “sea level equivalent” (A). (A) Ice volume stack from Spratt, R. M., and Lisiecki (2016). Climate of the Past 12, 1079–1092. (B) Core MD97-2142 from Ren, H., Sigman, D. M., Martínez-García, A., Anderson, R. F., Chen, M.-T., Christina Ravelo, A., Straub, M., Wong, G. T. F., and Haug, G. H. (2017), PNAS 114, 6759–6766. (C) Core ODP 999A from Straub, M., Sigman, D. M., Ren, H., Martínez-García, A., Meckler, A. N., Hain, M. P., and Haug, G. H. (2013), Nature 501, 200–204. (D) Core ODP 1012 from Liu, Z., Altabet, M. A., and Herbert, T. D. (2008), Geochemistry Geophysic Geosystems 9(11), Q11006, doi:10.1029/ 2008GC002044. (E) Core ODP 1090 from Martínez-García, A., Sigman D. M., Ren, H., Anderson, R. F., Straub, M., Hodell, D. A., Jaccard, S. L., Eglinton, T. I., and Haug, G. H. (2014). Science 343, 1347–1350. (F) Core PS75/072-4 from Studer, A. S., Sigman, D. M., Martínez-García, A., Benz, V., Winckler, G., Kuhn, G., Esper, O., Lamy, F., Jaccard, S. L., Wacker, L., Oleynik, S., Gersonde, R., and Haug, G. H. (2015). Paleoceanography 30, 845–862. 278 Ocean Process Tracers | Nitrogen Isotopes in the Ocean However, studies of the N isotopes in the modern ocean make clear that the d15N of the organic matter produced in and exported from the surface ocean, which the paleoproxy records seek to reconstruct, may be responding to more than one process. For example, a given d15N record from the low-latitude ocean could record overlapping changes in (1) N2 fixation, (2) mean ocean nitrate d15N, and (3) the d15N elevation of nitrate in inflowing shallow subsurface waters by partial nitrate assimilation in subpolar regions or by denitrification suboxic zones. The d15N of global mean ocean nitrate, which is close to that of deep ocean nitrate, is a particularly challenging parameter to reconstruct for the past ocean. In low-latitude settings, the mid-depth and thermocline waters sit above the deep water, preventing the d15N of the organic matter produced in the surface ocean from directly recording the d15N of deep nitrate (Fig. 2). Deep water can upwell or mix directly to the surface in some high-latitude regions, the Southern Ocean in particular. However, in these regions, the degree of nitrate consumption by phytoplankton assimilation may vary over time, so that d15N changes in biomass produced in the surface cannot be unambiguously interpreted as reflecting only changes in deep ocean nitrate d15N. This uncertainty in the history of global ocean nitrate d15N, in turn, compromises our ability to interpret down-core changes in d15N at any given site. Concluding Remarks The study of the N isotopes in the ocean is young relative to that of the other light isotopes (e.g., carbon, oxygen, and sulfur), with much of the work to date developing the methods needed to measure different forms of oceanic N and establishing the isotope systematics of N cycle processes that are necessary to interpret observed patterns. Over the previous decades, the N isotopes have had perhaps their greatest impact on food web studies and in paleoceanographic work. In the case of the latter, this reflects the ability of the N isotopes to provide basic constraints on environmental conditions when there are few other indicators available. Recent and ongoing method development is greatly improving our ability to measure diverse N pools in the ocean. This is yielding a new generation of N isotope studies that are beginning to provide geochemical estimates for the rates and distributions of N fluxes in the modern ocean, complementing instantaneous “bottle” measurements of these fluxes as well as other geochemical approaches. Fundamental aspects of the oceanic N cycle remain poorly understood, and the N isotopes provide an important tool for their study. Further Reading Altabet MA and Francois R (1994) Sedimentary nitrogen isotopic ratio as a recorder for surface ocean nitrate utilization. Global Biogeochemical Cycles 8(1): 103–116. Brandes JA and Devol AH (2002) A global marine fixed nitrogen isotopic budget: Implications for Holocene nitrogen cycling. Global Biogeochemical Cycles 16: 67-1–67-14. Casciotti KL (2016) Nitrogen and oxygen isotopic studies of the marine nitrogen cycle. Annual Review of Marine Science 8: 379–407. Casciotti KL and McIlvin MR (2007) Isotopic analyses of nitrate and nitrite from reference mixtures and application to Eastern Tropical North Pacific waters. Marine Chemistry 107: 184–201. Chang CCY, Silva SR, Kendall C, Michalski G, Casciotti KL, and Wankel S (2004) Preparation and analysis of nitrogen-bearing compounds in water for stable isotope ratio measurement. In: deGroot PA (ed.) Handbook of stable isotope analytical techniques, Vol. 1, pp. 305–354. Amsterdam: Elsevier. DeVries T, Deutsch C, Rafter PA, and Primeau F (2013) Marine denitrification rates determined from a global 3-D inverse model. Biogeosciences 10: 2481–2496. Galbraith ED, Sigman DM, Robinson RS, and Pedersen TF (2008) Nitrogen in past marine environments. In: Capone DG, Bronk DA, Mulholland MR, and Carpenter EJ (eds.) Nitrogen in the marine environment, pp. 1277–1302. Amsterdam: Elsevier. Granger J, Sigman DM, Needoba JA, and Harrison PJ (2004) Coupled nitrogen and oxygen isotope fractionation of nitrate during assimilation by cultures of marine phytoplankton. Limnology and Oceanography 49(5): 1763–1773. Hannides CCS, Popp BN, Choy CA, and Drazen JC (2013) Midwater zooplankton and suspended particle dynamics in the North Pacific Subtropical Gyre: A stable isotope perspective. Limnology and Oceanography 58(6): 1931–1946. Lehmann MF, Sigman DM, McCorkle DC, Granger J, Hoffmann S, Cane G, and Brunelle BG (2007) The distribution of nitrate 15N/14N in marine sediments and the impact of benthic nitrogen loss on the isotopic composition of oceanic nitrate. Geochimica et Cosmochimica Acta 71: 5384–5404. Montoya JP (2008) Nitrogen stable isotopes in marine environments. In: Capone DG, Bronk DA, Mulholland MR, and Carpenter EJ (eds.) Nitrogen in the marine environment, pp. 1277–1302. Amsterdam: Elsevier. Ohkouchi N, Chikaraishi Y, Close HG, Fry B, Larsen T, Madigan DJ, McCarthy MD, McMahon KW, Nagata T, Naito YI, Ogawa NO, Popp BN, Steffan S, Takano Y, Tayasu I, Wyatt ASJ, Yamaguchi YT, and Yokoyama Y (2017) Advances in the application of amino acid nitrogen isotopic analysis in ecological and biogeochemical studies. Organic Geochemistry 113: 150–174. Ren H, Sigman DM, Meckler AN, Plessen B, Robinson RS, Rosenthal Y, and Haug GH (2009) Foraminiferal isotope evidence of reduced nitrogen fixation in the ice age Atlantic Ocean. Science 323: 244–248. Sigman DM, Altabet MA, McCorkle DC, Francois R, and Fischer G (2000) The d15N of nitrate in the Southern Ocean: Nitrogen cycling and circulation in the ocean interior. Journal of Geophysical Research 105(C8): 19599–19614. Sigman DM, Granger J, DiFiore PJ, Lehmann MF, Ho R, Cane G, and van Geen A (2005) Coupled nitrogen and oxygen isotope measurements of nitrate along the eastern North Pacific margin. Global Biogeochemical Cycles 19: GB4022, doi:10.1029/2005GB002458. Sigman DM, DiFiore PJ, Hain MP, Deutsch C, Wang Y, Karl DM, Knapp AN, Lehmann MF, and Pantoja S (2009) The dual isotopes of deep nitrate as a constraint on the cycle and budget of oceanic fixed nitrogen. Deep Sea Research Part I 56: 1419–1439.