The marine nitrogen cycle: recent discoveries, uncertainties and the potential relevance of climate change

(a). Global patterns in dissolved inorganic nitrogen concentrations

Nitrogen is an essential element for all life forms but, in sea water, nitrogen mostly occurs as inert dissolved N₂ gas (more than 95%) that is inaccessible to most species. The rest is reactive nitrogen (N_r), such as nitrate, ammonia and dissolved organic compounds. Nitrate concentrations in surface waters vary from almost zero in the tropics and subtropics up to several tens of micromoles per litre in the temperate and Arctic and Antarctic oceans (figure 1a).

The increase in deep-water nutrient concentration from the northern Atlantic towards the Indian and Pacific oceans (figure 1b) is a result of a global circulation pattern, called the oceanic ‘conveyor belt’ [19]. The circulation is driven by the excess salt in the northern Atlantic Ocean where surface waters sink to the bottom and travel southwards. This southward current partly supplies water to the Antarctic Circumpolar Current, and continues to the Indian and the Pacific Oceans [19] from whence the flow returns to the surface to travel back through the Indonesian Islands the Tasman Sea [20], and further into the Indian and Atlantic Oceans. The residence time of the Atlantic water is around 180 years, whereas the complete overturning of ocean deep waters takes approximately 1000 years [19].

Four major processes deliver nutrient-rich waters to the surface: upwelling, winter mixing, eddies and diffusion from below a permanent or seasonal thermocline (table 1). Coastal upwelling is especially strongly developed in eastern boundary currents along the west coasts off North and South America and Africa, where nutrient-rich water from approximately 200 m is brought to the surface (figure 1a). The upwelled water stimulates primary production and the export of organic material, with the consequence of oxygen deficiency in subsurface waters owing to remineralization. The recycling of organic matter at depth leads to accumulation of nutrients with nitrate concentrations between 20 and 50 µmol l⁻¹ (figure 1b). At high latitudes, deep wintertime convection mixes the waters down to a depth of several hundred metres so that seasonal interfaces are destroyed. Thus, nutrients from greater depth are redistributed into the surface layer to provide the nutrient reservoir for the spring bloom in the following year. In the open ocean, eddy fields are generated by baroclinic or barotropic instabilities and large-scale gyres by the divergence of the horizontal wind field. Within these, structure nutrients are supplied to the surface by Ekman pumping, stimulating blooms of phytoplankton [21,22]. To what extent these sites contribute to global primary production is difficult to estimate, because the occurrence of eddies is transient and varies greatly in time and location. For the Atlantic Ocean, roughly one-third of the total N flux may be delivered by eddy transport [23]. Phytoplankton living in surface waters drives the N cycle and often consumes all inorganic nutrients down to a depth where ambient light level is 1–0.1% of the surface irradiation. At these depths, turbulent diffusion is another important source for primary production [24].

Table 1.

Summary of major nitrogen nutrient sources to the ocean.

source of nutrient	major sites of occurrence	example of sites where process is of significance	type of nutrient	timing
convective overturning	high latitudes	North Atlantic, Greenland and Norwegian Seas		autumn, winter
wind induced and Ekman upwelling	eastern boundary upwelling systems	off Oregon/California/Mexico, Peru/Chile, Iberian Peninsula/Mauretania, Namibia, NW Indian Ocean (Arabian Sea)	, DON	during upwelling season (depending on the direction of the wind fields)
eddy activity/gyres	everywhere	frontal systems; Gulf stream, Kuroshio, Polar front		sporadically/permanent
diffusion	thermoclines	oligotrophic ocean		permanent, depending on the gradient
rivers	continental shelves	Amazon, Mississippi, Yangtze, Mekong		permanent
atmospheric deposition	coastal, but globally relevant	oligotrophic tropical ocean		sporadically
shelf processes	on continental shelves	European shelf, Patagonian shelf		permanent
regeneration of particles	everywhere in the water column and in sediments	high latitude surface waters in summer, low latitude surface waters, benthic boundary layers		permanent

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Our understanding of ocean productivity regulation was significantly advanced when the concept of new and regenerated production was introduced by Dugdale & Goering [25]. They defined new production as that based on the input of nitrate from outside the euphotic zone (including nitrogen fixation), whereas regenerated production (based on ammonium) fuels a ‘microbial loop’ first introduced by Azam et al. [26] within a defined surface layer of the ocean. Based on Dugdale and Goerings's concept, only the amount of newly produced biomass can be quantitatively exported to the deep ocean without running down the primary production system [27]. However, this concept assumes that sources of new nitrogen are small, and that the ocean is in steady state, which often seems not to be the case [3,6,28]. Other potential problems with the Dugdale and Goering concept are bacterial uptake of dissolved inorganic nitrogen (DIN), assimilation of dissolved organic nitrogen (DON) and nitrification in the euphotic zone [29,30]. Of the global net primary production in the oceans of 1900 Tg C yr⁻¹ [31], 320 Tg yr⁻¹ should be N, using a C : N ratio of 6. To these, about 140 Tg N yr⁻¹ are added by nitrogen fixation [32], whereas global losses may be around 400 Tg N yr⁻¹ [33]. A global nitrogen budget considering all natural and human sources and incorporating recent discoveries has not yet been put together, but qualitative attempts have been made [3,6,32,34].

Nitrogen in atmospheric deposition develops into a significant N source in the open ocean, where usually very little combined nitrogen occurs in surface waters [15]. To what extent these atmospheric nutrients support oceanic productivity, and decrease oxygen concentrations, remains unclear, but Dentener et al. [35] estimate a total nitrogen deposition of 46.2 Tg N yr⁻¹, whereas Duce et al. [15] derive a global estimate of 67 Tg N yr⁻¹.

(b). Nitrogen fixation

Nitrogen-fixing organisms are independent of combined N sources and able to use the dissolved N₂ gas, which has concentrations of over 400 µmol l⁻¹ in sea water. In order to be used as a source of nitrogen, dinitrogen has to be reduced to ammonium for which specific enzymes are required [36]. Moreover, nitrogen-fixing species rely on phosphate and require iron in much larger quantities than other phytoplankton [37]. Therefore, co-limitation of Fe and P occurs [38] and may even regulate the global occurrence of nitrogen-fixing species. Because the Fe input from land is lower in the Pacific than the Atlantic Ocean [39], the tropical Atlantic Ocean may experience higher overall N fixation than the Pacific [40]. But recent evidence suggests higher rates in the Pacific Ocean than previously assumed, possibly higher than in the Atlantic Ocean [28,41].

Nitrogen fixation in the ocean has long been focused on the large bloom-forming genus Trichodesmium [6], until smaller unicellular species were discovered and their potential fixation activity demonstrated [7,42,43]. These are heterotrophic nitrogen fixers, including the photoheterotroph group-A (UCYN-A), which lacks the photosystem II [44], and Richelia intracellularis, a symbiont that fixes nitrogen in diatoms cells and has been found in tropical river plumes [45,46].

Nitrogen fixation may also be an important source of nitrogen to support biological production in deep-sea environments. Mehta et al. [47] described a diverse group of organisms possessing nifH genes, and therefore potentially capable of fixing N₂, from hydrothermal vent sites of the Juan de Fuca Ridge system. Other archaeal nifH sequences, many of uncertain phylogeny, have been linked to subsurface circulation through hydrothermal vent systems [48] and to methane cold seeps [9,49–51] and a mud volcano [52]. Further work at a cold seep field has revealed that methanotrophic archaea form syntrophic associations with sulfur-reducing bacteria, and that these consortia actively fix N₂ [9,50,51]. The overall contribution of these benthic diazotrophs to the oceanic nitrogen budget remains poorly known, but may be significant.

Nitrogen fixation rates based on direct measurements are sparse, and the variability is extremely high [6]. Therefore, global estimates of total marine nitrogen fixation are often based on geochemical factors such as unbalanced N : P ratios (N* method) which are assumed to support nitrogen fixation. Recently, Groszkopf et al. [53] reported a global marine N-fixation rate of 177 Tg N yr⁻¹ based on extrapolation of direct measurement, and suggest that N₂ fixers other than Trichodesmium are quite important outside the tropics. From global nutrient distributions and an ocean circulation model, Deutsch et al. [32] calculated a nitrogen fixation rate of 140 Tg N yr⁻¹, which agrees well with other estimates in the range of 100–200 Tg N yr⁻¹ [6]. It seems that lower estimates mainly result from model exercises where additional factors limiting N fixation such as iron have been implemented [54,55]. The global N-fixation rate of 140 Tg N yr⁻¹ has been largely unquestioned and is widely used [3].

(c). The role of dissolved organic nitrogen in nitrogen cycling

Dissolved organic nitrogen is rather uniformly distributed in the water column of the open ocean, with slightly higher concentrations in the surface than at depth, but with DON increasing considerably towards coastal areas and in estuaries [56]. With a mean concentration of 5.8 ± 2 µmol l⁻¹, DON may potentially be more important than inorganic forms, because DON concentrations comprise between 18 per cent and 85 per cent of the total nitrogen pool in coastal and open ocean surface water, respectively, with particulate nitrogen being negligible [56]. The DON pool is not as inert as suggested by the relatively high and constant concentrations found in the oceans, but a small part of it is rather dynamic and consumed by phytoplankton and bacteria [57]. Furthermore, DON is mainly of autochthonous origin because it stems from direct release by phytoplankton and bacteria [58,59], egestion and excretion from micro- and meso-zooplankton [60], or viral lysis of bacterioplankton [61].

In coastal waters, rivers are a major source of allochthonous DON, and its composition, bioavailability and quantities may vary with land use [62]. Another allochthonous source relevant also for the open ocean is the DON in atmospheric deposition [63,64]. DON is actively channelled into cells via membrane transport systems [65] and seems to be a quite active component of coastal nitrogen cycling [66,67]. Better understanding of the dynamic of DON is essential to quantify its role in the nitrogen and carbon cycles and how these will respond to anthropogenic perturbations and global change.

(d). Stoichiometry of C : N : P in the ocean

The close similarity between nitrogen and phosphorous ratios in plankton and in deep-water nutrients was first noted by Redfield [68], who suggested that life in the ocean adjusts the nutrients according to its requirements. Today, the perception is, rather, vice versa, that life has adjusted to the oceanic ratios. The C : N : P ratio of 106 : 16 : 1 on a molar basis is still a fundamental concept in marine sciences and mirrors the metabolic demands of an average living cell [69]. This ‘Redfield ratio’ allows linking elemental cycles and has been widely used in ecosystem and element flux models [32]. The cellular stoichiometry is important for the regulation of organic C and N cycling and also for the linkage between single-cell activity and ecosystem function. Heterotrophic organisms maintain low nitrate concentrations in the water when organic C : N ratios match the stoichiometric demands, but as soon as organic carbon becomes limiting, denitrification decreases, and nitrification is enhanced [70]. It may be quite important to focus nutrient management scenarios to maintain balanced elemental ratios because nitrogen supply in excess of the demand leads to significant enrichment of N_r [71].

The ultimately limiting element for ocean productivity had been debated over decades, and a consensus seems to be that nitrogen is limiting on short time-scales, whereas on geological time scales it seems to be phosphorus [72,73]. Both these views are based on input and loss rates and the turnover of the elements, which all may vary over time [74]. However, human perturbations affecting the input or loss of substances ultimately impact the global elemental cycles and the productivity.

(a). Nitrous oxide production, consumption and release from the ocean

The marine pathways of N₂O and the quantification of its oceanic emissions have received increased attention during recent decades [82,83]. N₂O acts as a strong greenhouse gas and in the stratosphere it is the precursor of ozone-depleting nitric oxide (NO) radicals [84,85]. Because of the ongoing decline of chlorofluorocarbons and the continuous increase of N₂O in the atmosphere, the contributions of N₂O to both the greenhouse effect and ozone depletion will be even more pronounced in the twenty-first century [84,86]. The oceans, including coastal areas such as continental shelves, estuaries and upwelling areas are a major source of N₂O and contribute about 30 per cent (5.5 Tg N yr⁻¹) to the atmospheric N₂O budget [84]. Oceanic N₂O is produced as a by-product during bacterial as well as archaeal nitrification; however, nitrifying archaea dominate N₂O production [87,88]. N₂O occurs also as an intermediate during denitrification. It is produced by the reduction of and, in turn, can be further reduced to N₂. Nitrification is the dominating N₂O production process, whereas denitrification contributes only 7–35% to the overall N₂O water column budget in the ocean [89,90]. The importance of other microbial processes, such as dissimilatory nitrate reduction to ammonia and anammox, for oceanic, N₂O production is largely unknown [83].

The amount of N₂O produced during both nitrification and denitrification strongly depends on the prevailing dissolved oxygen (O₂) concentrations and is significantly enhanced under low (i.e. suboxic) O₂ conditions [2,91]. N₂O is usually not detectable in anoxic waters because of its reduction to N₂ during denitrification. Thus, significantly enhanced N₂O concentrations are generally found at oxic/suboxic boundaries in the oceans [2,83]. The strong O₂ sensitivity of N₂O production is also observed in coastal systems that are characterized by seasonal shifts in the O₂ regime [92]. A biological source of N₂O in the well-oxygenated, mixed layer/euphotic zone seems to be unlikely (see discussion in Freing et al. [90]). A recent study of both the N₂O air/sea fluxes and N₂O diapycnal fluxes into the mixed layer in the coastal upwelling off Mauritania, northwest Africa indicated that surfactants may have a dampening effect on air–sea exchanges of N₂O [93]. However, for a reassessment of the global coastal N₂O emissions, further laboratory and field studies of the effect of surfactants on the air/sea flux of N₂O are needed.

Global maps of N₂O in the surface ocean [94,95] show supersaturation of N₂O in coastal and equatorial upwelling regions as well as N₂O anomalies close to zero (i.e. near equilibrium) in large parts of the open ocean. Since the studies of Nevison et al. [94] and Suntharalingam & Sarmiento [95], the amount of available N₂O data has been steadily increasing. Therefore, the project ‘Marine Methane and Nitrous Oxide’ database has been launched with the aim to collect and archive N₂O datasets and to provide actual fields of surface N₂O for emission estimates [96].

The observed ongoing deoxygenation of the open ocean [97] can be translated into an additional N₂O accumulation of less than 6 per cent [83]. In contrast to the open ocean, Naqvi et al. [92] cautioned that global N₂O emissions from shallow hypoxic/anoxic coastal systems might increase significantly in the future. It seems realistic to expect that the N₂O emissions from shallow hypoxic/anoxic coastal systems will increase in the near future due to increasing nutrient inputs (caused by the ongoing industrialization and intensification of agricultural activities), whereas future N₂O emissions from the open ocean owing to decreasing oxygen in mid water depths seem to be of minor significance.

(a). Proximal coastal systems

Each year, rivers transport about 40–66 Tg N yr⁻¹ to coastal ecosystems, 40 per cent in the form of DIN (mainly nitrate), 40 per cent in the form of particulate nitrogen and the remaining 20 per cent as dissolved organic nitrogen [62]. The variability in the export is impacted by anthropogenic nitrogen input, whereas DON export varies between 10 and 25 per cent of total N [62]. Submarine groundwater discharge delivers about 4 Tg N yr⁻¹ to near-shore ecosystems. Further inputs of nitrogen to the proximate coastal ocean include nitrogen fixation in the water column, sediments and vegetated ecosystems totalling about 15 Tg N yr⁻¹ and atmospheric deposition (about 1 Tg N yr⁻¹). Losses of fixed nitrogen include export to the continental shelf (38 Tg N yr⁻¹), emission of N₂O (0.5 Tg N yr⁻¹), denitrification (4–8 Tg N yr⁻¹), fish landings (3.7 Tg N yr⁻¹) and burial (22 Tg N yr⁻¹; figure 2). This burial occurs mainly in vegetated systems such as seagrass meadows and mangroves [98], because of trapping of particulate nitrogen and assimilation of dissolved inorganic and organic nitrogen [67]. The present-day and future nitrogen cycling in proximal coastal systems differs substantially from that before the Anthropocene. Eutrophication has resulted in community shifts such as a loss of submerged vegetation and proliferation of macroalgae and phytoplankton. This will eventually lead to lower rates of nitrogen burial. Denitrification may be higher or lower depending on conditions [99], whereas N₂O emission is higher because of extensive oxic–anoxic interfaces with high rates of nitrification and denitrification.

(b). Distal coastal systems

The fixed nitrogen budget of continental shelf systems is primarily governed by exchanges with the open ocean (figure 2). Globally, continental shelves receive about 38 Tg N yr⁻¹ exported from the proximal coastal zone, 7.5 Tg N yr⁻¹ from atmospheric deposition (almost all anthropogenic) and a small contribution by nitrogen-fixing organisms (about 2 Tg N yr⁻¹). Continental shelf systems lose nitrogen via denitrification (100–250 Tg N yr⁻¹); the uncertainty, however, is high [100]. Because the global estimate of 400 Tg N yr⁻¹ removal per year exists, about 200–300 Tg N yr⁻¹ remain for the open ocean (figure 2). Recent estimates suggested five times higher denitrification rates in permeable sediments [101], so the number may need an upward revision in the future. Most nitrogen entering the continental shelf comes from nitrate exchanges with the open ocean (450–600 Tg N yr⁻¹). The majority of the nitrogen imported from subsurface oceanic waters is returned in the form of DON and particulate organic nitrogen (PON) to open ocean surface waters (about 390 Tg N yr⁻¹). This organic nitrogen exported from coastal systems supports the reported heterotrophy of oceanic gyre ecosystems and constitutes a source of fixed nitrogen for micro-organisms. A budget of these numbers for the three systems (figure 2) reveals a relatively low uncertainty in the numbers around 10 per cent that supports the reliability of the estimates.