Nutrient conversion in the marine environment

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Although there is uncertainty about the estimates of nutrient conversion in the marine environment, it is widely believed that the underlying biogeochemical processes largely take place in the sediment of the continental shelf. This article provides a brief introduction to the processes involved in the conversion of the main nutrients: nitrogen, phosphorus and silicon.


Nutrient sources

Nutrients in coastal environments originate from multiple sources, including rivers, atmospheric deposition, groundwater discharge, and in situ biological fixation (see also What causes eutrophication?). Rivers are the primary transport pathway of nutrients to coastal oceans. Riverine inputs of nitrogen (N), mainly as nitrate (NO3−) and ammonia (NH3)/ammonium (NH4+), and phosphorus (P), mainly as orthophosphate (PO43−), approximately doubled between 1960 and 1990[1]. This increase is largely attributable to the expanded use of synthetic fertilizers in agriculture. Additional anthropogenic sources include wastewater discharge from urban sewer systems[2] and industrial effluents. Natural processes such as rock weathering[3] and terrestrial nitrogen fixation[4] also contribute to riverine nitrogen export.

Around 2015, global riverine nitrogen discharge to coastal waters was estimated at 35–50 Tg N yr−1[5][6][7][8][9]. Riverine discharge of reactive phosphorus is estimated at 2–5 Tg P yr−1[10][1][5][9] (1 Tg = 1012 g).

Submarine groundwater discharge may contribute nutrient fluxes comparable to or exceeding riverine inputs globally, although estimates remain uncertain[11]. Atmospheric deposition contributes approximately 8 Tg N yr−1 to continental shelves and 40–50 Tg N yr−1 to the global ocean[6]. Regional differences are substantial. For example, atmospheric deposition accounts for about 30% of total land-based nitrogen inputs to the North Sea and about 50% to the Baltic Sea[12]. The N:P ratio of atmospheric deposition can be very high; in the North Sea it was estimated at 503:1, far exceeding the Redfield ratio of 16:1 typical of phytoplankton[13].

In addition to external inputs from land, bioavailable nitrogen is produced in situ through biological nitrogen fixation (diazotrophy). Specialized diazotrophic cyanobacteria, such as the photoautotroph Trichodesmium and symbiotic unicellular cyanobacteria (e.g. UCYN-A), convert dissolved dinitrogen gas (N2) into ammonium (NH4+) using the enzyme nitrogenase[14][15]. This process occurs predominantly in warm, well-lit, oligotrophic ocean regions. Global marine nitrogen fixation is estimated at 70–170 Tg N yr−1[16][17][18]. Additional nitrogen fixation occurs in aphotic ocean regions by bacteria and archaea, contributing an estimated 13–134 Tg N yr−1[19]. Coastal and benthic nitrogen fixation adds approximately 15 Tg N yr−1[17].

In addition to nitrogen and phosphorus, silicon (Si) is an essential nutrient for silicifying organisms such as diatoms, silicoflagellates, certain radiolarians, rhizarians, and siliceous sponges. Dissolved Si (dSi), mainly as undissociated monomeric silicic acid, Si(OH)4, is the only Si compound available for uptake by marine organisms. Rivers are the main source of dissolved silicon to the ocean, followed by submarine groundwater, mineral dissolution, hydrothermal inputs, aeolian dust, and glacial weathering. A large reservoir of dissolved silicon exists in the deep ocean, which becomes available to surface ecosystems through upwelling[20].


Nutrient cycling

Fig.1. Bioturbated seabed.

Nutrient cycling is a process in which marine microorganisms play a crucial role. The key steps include the uptake of nutrients by phytoplankton to produce organic matter and the subsequent release of these nutrients during respiration and mineralization of organic matter by bacteria, making them available again for biological uptake (see Plankton bloom). Coastal shelf seas are zones of intense nutrient cycling, which enhances primary productivity[21].

Shelf seas are estimated to account for up to 80% of global benthic mineralization, despite covering only about 7% of the seafloor[22]. Most mineralization occurs after detrital plankton settles on the seabed and is decomposed under aerobic or anaerobic conditions.

In coarse-grained sediments, organic matter is rapidly mineralized because oxygen penetrates deeply into the sediment due to its high permeability[23]. As a result, oxygen consumption rates are high and organic carbon storage is low[24]. Mineralization is further enhanced by bioturbating benthic macrofauna, which increase sediment mixing and oxygen penetration[25] (Fig. 1).

In fine-grained cohesive sediments, oxygen penetration is limited; anaerobic mineralization pathways therefore become relatively more important, and organic carbon accumulates to higher levels[26].

In deeper parts of the continental shelf and the open ocean, less organic material reaches the seabed because a larger fraction is decomposed in the water column before settling[27]. See also Ocean carbon sink.


Nutrient transformation

Nitrogen

Nitrogen (N) species in aquatic environments include dissolved inorganic forms (nitrate NO3−, nitrite NO2−, ammonium NH4+), dissolved organic nitrogen, and particulate organic nitrogen[28]. Nitrogen is removed from aquatic systems by sedimentation and burial, and most importantly by conversion to atmospheric gases through denitrification coupled to organic matter decomposition.

Denitrification is the microbial reduction of nitrate to quasi-inert gases, mainly dinitrogen (N2) and nitrous oxide (N2O), during the oxidation of organic matter. It is a heterotrophic process that occurs under low-oxygen conditions (typically ≤0.2 mg O2 L−1)[29][21]. Denitrifying bacteria are widespread; the main controlling factors are nitrate or nitrite availability, low oxygen concentrations, and the presence of labile organic carbon. This organic carbon, including extracellular polymeric substances (EPS) produced by marine microorganisms, serves as an electron donor in the denitrification process.

Denitrification generally depends on prior nitrification, in which ammonium is oxidized to nitrite and nitrate by autotrophic bacteria and archaea under aerobic conditions. In the presence of comammox bacteria, ammonium can be directly oxidized to nitrate (COMplete AMMonia OXidation). Because nitrification requires oxygen and denitrification requires low oxygen conditions, their coupling occurs primarily in environments with sharp oxygen gradients. Muddy shelf-sea sediments provide such conditions[30][29]. Nitrate produced in the aerobic surface layer diffuses into deeper suboxic layers, where denitrification takes place.

Nitrogen can also be removed in the absence of organic matter through anaerobic ammonium oxidation. This microbial process, known as anammox, converts ammonium and nitrite directly into dinitrogen gas (N2) and water under anoxic conditions[31]. Biological activity of benthic macrofauna can substantially extend the depth of the oxic–suboxic interface through bioturbation and bioirrigation, thereby enhancing coupled nitrification–denitrification and increasing overall nitrogen removal[32].

Denitrifying bacteria compete with microorganisms performing dissimilatory nitrate reduction to ammonium (DNRA). Unlike denitrification, DNRA retains nitrogen in the ecosystem by converting nitrate into ammonium, which remains available for biological uptake. DNRA is common in low-oxygen environments such as hypoxic sediments and oxygen minimum zones. Ammonium produced by DNRA can subsequently be converted to dinitrogen gas through anammox.

Globally, most denitrification occurs in shelf-sea sediments (Fig. 2), where nitrogen removal is estimated at 200–300 Tg N yr−1[29]. Additional denitrification occurs in oxygen-deficient regions of the open ocean, particularly in the Eastern Tropical North Pacific, Eastern Tropical South Pacific, and the Arabian Sea, where approximately 50–100 Tg N yr−1 is removed[33][29].

Denitrification also occurs in estuaries. Significant denitrification has been observed where a strong estuarine turbidity maximum is present[34]; it is surmised that aerobic degradation of suspended organic particles creates hypoxic microniches for the growth of denitrifying bacteria. Important denitrification has also been reported in benthic intertidal ecosystems, both in the freshwater and salt water tidal reaches[35], with possibly an important role for microalgae[36]. A field study in a tropical estuary (Northern Australia) revealed intertidal mudflats per unit area to be hotspots of N removal, with higher denitrification efficiency than other benthic habitats, including mangrove and saltmarsh[37].

Denitrification is enhanced with increasing temperatures, at least partly because higher temperatures favor anoxic conditions[38]. There is also evidence that denitrifying microbial communities are sensitive to salinity fluctuations[39].


Fig. 2. Illustration of the influence of hypoxia on nutrient conversion. OM=organic matter. Redrawn after Dai et al. (2023[40]).


Phosphorus

Phosphorus (P) discharged by rivers into coastal waters occurs in dissolved and particulate forms, including dissolved inorganic phosphorus (DIP), dissolved organic phosphorus (DOP), particulate inorganic phosphorus (PIP), and particulate organic phosphorus (POP). Much of the non-reactive PIP, such as apatite, is deposited on continental shelves and does not reach the open ocean[41].

A significant fraction of riverine phosphate is adsorbed onto clay particles and associated with iron and manganese oxides or oxyhydroxides. These particles are often retained in estuaries, but increasing salinity promotes desorption, releasing phosphate into coastal waters[42][43]. It has been estimated that the amount of phosphate released through desorption from suspended particles may be two to five times greater than the dissolved phosphate directly delivered by rivers[41].

Hypoxic conditions in coastal waters promote the release of dissolved inorganic phosphorus from sediments. Under low-oxygen conditions, iron oxides that bind phosphate are reduced, releasing phosphate into pore water and subsequently into the overlying water column[44]. This internal phosphorus recycling can sustain or enhance algal blooms (Fig. 2).

Particulate phosphorus deposited in sediments may also be released into pore waters during diagenesis, contributing to benthic phosphorus fluxes to bottom waters[41]. Dissolved inorganic phosphorus, mainly in the form of orthophosphate (PO43−), is assimilated by phytoplankton and incorporated into organic matter, forming DOP and POP. This organic phosphorus is returned to the dissolved inorganic pool through microbial mineralization during decomposition. Most DOP is mineralized in surface waters, whereas the fraction transported to deeper ocean layers may persist for thousands of years due to slow turnover rates. Permanent removal of phosphorus from the ocean occurs primarily through burial in marine sediments, estimated at approximately 3–10 Tg P yr−1[45].

Phosphorus may also be removed through microbial reduction of phosphate to gaseous phosphine (PH3), although this pathway is poorly quantified and its global significance remains uncertain[46][28].


Silicon

Silicon (Si) in the marine environment occurs mainly as dissolved silicon (dSi), primarily in the form of silicic acid Si(OH)4, and as particulate silicon in the form of biogenic silica (bSiO2, also called opal). Biogenic silica consists of amorphous silica contained in living organisms and in biogenic detritus in the water column, soils, and sediments. The main transformation processes are the biological uptake of dissolved silicon and its biomineralization into biogenic silica, and the subsequent dissolution of biogenic silica back into dissolved silicon.

Over long time scales, biogenic silica undergoes chemical and mineralogical transformations during diagenesis[47]. These transformations may include conversion of opaline silica into authigenic aluminosilicate minerals such as clays[48].

Diatoms are the dominant producers of biogenic silica in the marine environment. Radiolarians, siliceous sponges, and chrysophytes can also contribute locally to biogenic silica production[49]. Global marine biogenic silica production is estimated at 7000–8000 Tg Si yr−1[2]. Most biogenic silica produced in surface waters dissolves in the upper ocean or during sinking, with only about one third reaching the seafloor[50]. Much of the deposited silica is recycled at the sediment–water interface, and only a small fraction, approximately 250 Tg Si yr−1, is ultimately buried in marine sediments. The total oceanic dissolved silicon inventory is estimated at about 3.36 × 106 Tg Si, most of which resides in the deep ocean[2].

The primary external source of dissolved silicon to the ocean is river discharge, derived from continental weathering. Total dissolved silicon input to the ocean, including contributions from groundwater, hydrothermal systems, and atmospheric dust, is estimated at 340–490 Tg Si yr−1[2]. Losses from the marine silicon pool occur mainly through sediment burial and reverse weathering, in which dissolved or biogenic silica is incorporated into authigenic aluminosilicate minerals, as well as through accumulation in siliceous sponge skeletons[2].

Although these fluxes are small relative to the large oceanic dissolved silicon reservoir, regional silicon availability is sensitive to changes in supply and removal. Human activities such as dam construction reduce silicon delivery to coastal waters by trapping biogenic and particulate silica in reservoirs. This effect has already altered silicon cycling in coastal zones downstream of major dammed rivers such as the Nile and Danube[20]. Reduced silicon availability can favor non-siliceous phytoplankton over diatoms, potentially altering marine food webs and biogeochemical cycling[51].

Climate change is also expected to influence the marine silicon cycle. Increased stratification of the ocean surface reduces vertical transport of dissolved silicon from deep waters, potentially limiting diatom growth. Conversely, melting sea ice in polar regions may locally increase dissolved silicon availability. Model projections suggest a global decline in diatom biomass over the coming century, except in parts of the Southern Ocean[52][53].


Trace metals

Very small concentrations of dissolved metals, known as trace metals (including iron Fe, manganese Mn, zinc Zn, copper Cu, cobalt Co, nickel Ni, and cadmium Cd), play essential roles in marine biogeochemical cycles. Trace metals are supplied to continental shelves through rivers and estuaries, where a substantial fraction is removed by adsorption onto fine sediment particles. In contrast, iron and manganese are delivered to the open ocean primarily through atmospheric deposition of mineral dust. Additional sources include hydrothermal vents and sediment release.

Trace metals occur in various chemical forms, and their biological availability is regulated by photochemical reactions and microbial processes that modify their oxidation state and complexation with organic ligands[54]. Rapid recycling of trace metals takes place in surface waters, where dissolved concentrations are extremely low due to efficient biological uptake and regeneration.

Trace metals are essential components of metalloenzymes that regulate key metabolic processes and serve structural roles in proteins. Although required only in trace amounts, both deficiency and excess can be harmful: insufficient availability limits biological productivity, while elevated concentrations may be toxic. Marine organisms have evolved to function within narrow trace metal concentration ranges, making marine ecosystems sensitive to environmental changes that alter trace metal availability. Ocean acidification may influence trace metal speciation and bioavailability, although the mechanisms remain incompletely understood (see Ocean acidification#Bioavailability of trace metals).

Trace metals strongly influence primary production and nutrient cycling. For example, phytoplankton growth can be limited by iron availability in large ocean regions. Iron is also essential for nitrogen fixation by diazotrophic cyanobacteria such as Trichodesmium. Limited trace metal availability can therefore constrain the marine nitrogen cycle, since many nitrogen transformation processes rely on metalloenzymes containing iron, copper, or molybdenum. Trace metals are also involved in carbon fixation (requiring Fe and Mn), organic matter remineralization (Fe, Zn), methane oxidation (Cu), calcification (Zn, Co), and silica biomineralization in some organisms (Zn, Cd, Se)[54].


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References

  1. ↑ 1.0 1.1 Beusen, A. H. W., Bouwman, A. F., Van Beek, L. P. H., Mogollón, J. M. and Middelburg, J.J. 2015. Global riverine N and P transport to ocean increased during the twentieth century despite increased retention along the aquatic continuum. Biogeosciences Discuss. 12: 20123–20148
  2. ↑ 2.0 2.1 2.2 2.3 2.4 Tuholske, C., Halpern, B.S., Blasco, G., Villasenor, J.C., Frazier, M. and Caylor, K. 2021. Mapping global inputs and impacts from human sewage in coastal ecosystems. PLoS ONE 16(11), e0258898 Cite error: Invalid <ref> tag; name "T21" defined multiple times with different content
  3. ↑ Houlton, B.Z., Morford, S.L. and Dahlgren, R.A. 2018. Convergent evidence for widespread rock nitrogen sources in Earth’s surface environment. Science 360: 58–62
  4. ↑ Davies‐Barnard, T. and Friedlingstein, P. 2020. The global distribution of biological nitrogen fixation in terrestrial natural ecosystems. Global Biogeochemical Cycles 34, e2019GB006387
  5. ↑ 5.0 5.1 Beusen, A.H.W. and Bouwman, A.F. 2022. Future projections of river nutrient export to the global coastal ocean show persisting nitrogen and phosphorus distortion. Front. Water 4, 893585
  6. ↑ 6.0 6.1 Malone, T.C. and Newton, A. 2020. The globalization of cultural eutrophication in the coastal ocean: Causes and consequences. Front. Mar. Sci. 7: 670
  7. ↑ Yamamoto, A., Hajima, T., Yamazaki, D., Aita, M.N., Ito, A., Kawamiya, M. 2022. Competing and accelerating effects of anthropogenic nutrient inputs on climate-driven changes in ocean carbon and oxygen cycles. Science Advances 8, eabl9207
  8. ↑ Tivig, M., Keller, D.P. and Oschlies, A. 2021. Riverine nitrogen supply to the global ocean and its limited impact on global marine primary production: a feedback study using an Earth system model. Biogeosciences 18: 5327–5350
  9. ↑ 9.0 9.1 Micella, I., Kroeze, C., Bak, M.P. and Strokal, M. 2024. Causes of coastal waters pollution with nutrients, chemicals and plastics worldwide. Marine Pollution Bulletin 198, 115902
  10. ↑ Seitzinger, S.P., Mayorga, E., Bouwman, A.F., Kroeze, C., Beusen, A.H.W., Billen, G., Van Drecht, G., Dumont, E., Fekete, B.M., Garnier, J. and Harrison, J.A. 2010. Global river nutrient export: A scenario analysis of past and future trends. Global Biogeochemical Cycles 24, GB0A08
  11. ↑ Wilson, S.J. et al. 2024. Global subterranean estuaries modify groundwater nutrient loading to the ocean. Limnology and Oceanography Letters 9, 411–422
  12. ↑ North Sea Task Force 1993. North Sea Quality Status Report. Oslo and Paris Commissions.
  13. ↑ Rendell, A. R., Ottley, C. J., Jickells, T. D. and Harrison, R. M. 1993. Atmospheric nitrogen inputs to the North Sea. Tellus 45B: 53–63
  14. ↑ Hallstrøm, S., Benavides, M., Salamon, E.R., Arístegui, J. and Riemann, L. 2022. Activity and distribution of diazotrophic communities across the Cape Verde Frontal Zone in the Northeast Atlantic Ocean. Biogeochemistry 160: 49–67
  15. ↑ Martínez-Pérez, C., Mohr, W., Löscher, C. et al. 2016. The unicellular diazotrophic symbiont UCYN-A is a key player in the marine nitrogen cycle. Nat. Microbiol. 1, 16163
  16. ↑ Galloway, J. N. et al. 2004. Nitrogen cycles: past, present, and future. Biogeochemistry 70: 153–226
  17. ↑ 17.0 17.1 Voss, M., Bange, H.W., Dippner, J.W., Middelburg, J.J., Montoya, J.P. and Ward, B. 2013. The marine nitrogen cycle: recent discoveries and uncertainties. Phil. Trans. R. Soc. B 368: 20130121
  18. ↑ Tang, W., Li, Z. and Cassar, N. 2019. Machine learning estimates of global marine nitrogen fixation. J. Geophys. Res. Biogeosci. 124: 717–730
  19. ↑ Benavides, M., Bonnet, S., Berman-Frank, I. and Riemann, L. 2018. Deep ocean nitrogen fixation. Front. Mar. Sci. 5: 108
  20. ↑ 20.0 20.1 Ittekkot, V., Humborg, C. and Schaefer, P. 2000. Hydrological alterations and marine biogeochemistry: a silicate issue? BioScience 50: 776–782
  21. ↑ 21.0 21.1 De Borger, E., Braeckman, U. and Soetaert, K. 2021. Rapid organic matter cycling in North Sea sediments. Continental Shelf Research 214: 104327
  22. ↑ Wollast, R. 1998. Evaluation and comparison of the global carbon cycle in the coastal zone and in the open ocean. In: Brink, K.H., Robinson, A. (Eds.), The Sea. John Wiley & Sons, Inc., Hoboken, New Jersey, pp. 213–252
  23. ↑ Huettel, M., Berg, P. and Kostka, J.E. 2014. Benthic exchange and biogeochemical cycling in permeable sediments. Ann. Rev. Mar. Sci. 6: 23–51
  24. ↑ Braeckman, U., Foshtomi, M.Y., Van Gansbeke, D., Meysman, F., Soetaert, K., Vincx, M. and Vanaverbeke, J. 2014. Variable importance of macrofaunal functional biodiversity for biogeochemical cycling in temperate coastal sediments. Ecosystems 17: 720–737
  25. ↑ Toussaint, E., De Borger, E., Braeckman, U., De Backer, A., Soetaert, K. and Vanaverbeke, J. 2021. Faunal and environmental drivers of carbon and nitrogen cycling along a permeability gradient in shallow North Sea sediments. Science of the Total Environment 767: 144994
  26. ↑ Canfield, D., Jørgensen, B., Fossing, H., Glud, R., Gundersen, J., Ramsing, N., Thamdrup, B., Hansen, J., Nielsen, L. and Hall, P.O. 1993. Pathways of organic carbon oxidation in three continental margin sediments. Mar. Geol. 113: 27–40
  27. ↑ Middelburg, J.J., Soetaert, K., Herman, P.M.J. and Heip, C.H.R. 1996. Denitrification in marine sediments: a model study. Global Biogeochem. Cycles 10: 661–673
  28. ↑ 28.0 28.1 Tappin, A.D. 2000. An examination of the fluxes of nitrogen and phosphorus in temperate and tropical estuaries: current estimates and uncertainties. Estuarine, Coastal and Shelf Science 55: 885–901
  29. ↑ 29.0 29.1 29.2 29.3 Seitzinger, S., Harrison, J.A., Böhlke, J.K., Bouwman, A.F., Lowrance, R., Peterson, B., Tobias, C. and Van Drecht, G. 2006. Denitrification across landscapes and waterscapes: A synthesis. Ecological Applications 16: 2064–2090
  30. ↑ Malcolm, S.J. and Sivyer, D.B. 1997. Nutrient recycling in intertidal sediments. In: Jickells, T. and Rae, J.E. (Eds.), Biogeochemistry of Intertidal Sediments. Cambridge University Press, pp. 59–83
  31. ↑ Burgin, A. and Hamilton, S. 2007. Have we overemphasized the role of denitrification in aquatic ecosystems? A review of nitrate removal pathways. Frontiers in Ecology and the Environment 5: 89–96
  32. ↑ Rysgaard, S. B., Christensen, P.B. and Nielsen, L. P. 1995. Seasonal variation in nitrification and denitrification in estuarine sediment colonized by benthic microalgae and bioturbating fauna. Marine Ecology Progress Series 126: 111–121
  33. ↑ Deutsch, C., Gruber, N., Key, R. M. and Sarmiento, J. L. 2001. Denitrification and N2 fixation in the Pacific Ocean. Global Biogeochemical Cycles 15: 483–506
  34. ↑ Zheng, Y., Hou, L., Zhang, Z., Ge, J., Li, M., Yin, G., Han, P., Dong, H., Liang, X., Gao, J., Gao, D. and Liu, M. 2021. Overlooked contribution of water column to nitrogen removal in estuarine turbidity maximum zone (TMZ). Science of the Total Environment 788: 147736
  35. ↑ O'Connor, J.A., Erler, D.V., Ferguson, A. and Mahler, D.Y. 2022. The tidal freshwater river zone: Physical properties and biogeochemical contribution to estuarine hypoxia and acidification - The “hydrologic switch”. Estuarine, Coastal and Shelf Science 268, 107786
  36. ↑ Laverman, A.M., Morelle, J., Roose-Amsaleg, C. and Pannard, A. 2021. Estuarine benthic nitrate reduction rates: Potential role of microalgae? Estuarine, Coastal and Shelf Science 257, 107394
  37. ↑ Fortune, J., Butler, E.C.V. and Gibb, K. 2023. Estuarine benthic habitats provide an important ecosystem service regulating the nitrogen cycle. Marine Environmental Research 190, 106121
  38. ↑ Veraart, A.J., de Klein, J.J. and Scheffer, M. 2011. Warming can boost denitrification disproportionately due to altered oxygen dynamics. PLoS One 6(3):e18508
  39. ↑ Marks, B.M., Chambers, L. and White, J.R. 2016. Effect of Fluctuating Salinity on Potential Denitrification in Coastal Wetland Soil and Sediments. Soil Sci. Soc. Am. J. 80: 516–526
  40. ↑ Dai, M., Zhao, Y., Cha,i F., Chen, M., Chen, N., Chen, Y., Cheng, D., Gan, J., Guan, D., Hong, Y., Huang, J., Lee, Y., Leung, K.M.Y., Lim, P.E., Lin, S., Lin, X., Liu, X., Liu, Z., Luo, Y-W., Meng, F., Sangmanee, C., Shen, Y., Uthaipan, K., Wan Talaat, W.I.A., Wan, X.S., Wang, C., Wang, D., Wang, G., Wang, S., Wang, Y., Wang, Y., Wang, Z., Wang, Z., Xu, Y., Yang, J-Y.T,, Yang, Y., Yasuhara, M., Yu, D., Yu, J., Yu, L., Zhang, Z. and Zhang, Z. 2023. Persistent eutrophication and hypoxia in the coastal ocean. Cambridge Prisms: Coastal Futures 1, e19, 1–28
  41. ↑ 41.0 41.1 41.2 Paytan, A. and McLaughlin, K. 2007. The oceanic phosphorus cycle. Chem. Rev. 107: 563–576
  42. ↑ Krom, M.D. and Berner, R.A. 1980. Adsorption of phosphate in anoxic marine sediments. Limnol. Oceanogr. 25: 797–806
  43. ↑ Frossard, E., Brossard, M., Hedley, M.J. and Metherell, A. 1995. Reactions controlling phosphorus cycling in soils. In: Tiessen, H. (Ed.), Phosphorus in the Global Environment. John Wiley & Sons, pp. 107–138
  44. ↑ Middelburg, J.J. and Levin, L.A. 2009. Coastal hypoxia and sediment biogeochemistry. Biogeosciences 6: 1273–1293
  45. ↑ Benitez-Nelson, C.R. 2000. The biogeochemical cycling of phosphorus in marine systems. Earth-Science Reviews 51: 109–135
  46. ↑ Gassman, G. 1994. Phosphine in the fluvial and marine hydrosphere. Marine Chemistry 45: 197–205
  47. ↑ Van Cappellen, P., Dixit, S. and van Beusekom, J. 2002. Biogenic silica dissolution in the oceans: reconciling experimental and field-based dissolution rates. Global Biogeochemical Cycles 16, 1075
  48. ↑ Michalopoulos, P., Aller, R.C. and Reeder, R.J. 2000. Conversion of diatoms to clays during early diagenesis in tropical continental shelf muds. Geology 28: 1095–1098
  49. ↑ Simpson, T.L. and Volcani, B.E. 1981. Silicon and siliceous structures in biological systems. Springer-Verlag, New York
  50. ↑ Tréguer, P. J. and De La Rocha, C. L. 2013. The world ocean silica cycle. Annu. Rev. Mar. Sci. 5: 477–501
  51. ↑ Humborg, C., Conley, D.J., Rahm, L., Wulff, F., Cociasu, A. and Ittekkot, V. 2000. Silicon retention in river basins: far-reaching effects on coastal marine ecosystems. Ambio 29: 45–50
  52. ↑ Bopp, L., Aumont, O., Cadule, P., Alvain, S. and Gehlen, M. 2005. Response of diatom distribution to global warming. Geophys. Res. Lett. 32
  53. ↑ Laufkötter, C. et al. 2015. Drivers and uncertainties of future global marine primary production. Biogeosciences 12: 6955–6984
  54. ↑ 54.0 54.1 Morel, F.M.M. and Price, N.M. 2003. The biogeochemical cycles of trace metals in the oceans. Science 300: 944–947


The main authors of this article are Job Dronkers, Pierre Regnier and Claudette Spiteri
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