Encyclopedia of Geochemistry

Living Edition
| Editors: William M. White

Ocean Salinity, Major Elements, and Thermohaline Circulation

  • Hein J. W. de BaarEmail author
  • Steven M. A. C. van Heuven
  • Rob Middag
Living reference work entry
DOI: https://doi.org/10.1007/978-3-319-39193-9_120-1


Salinity is the total amount of dissolved salts in seawater, generally reported in grams/kg (parts per thousand). Salinity is in turn dominated by only a few major elements in ionic form: Na+, Mg2+, Ca2+, K+, Sr2+, Cl, SO4 2−, Br, F, and HCO3 . Temperature together with salinity determines the density of seawater, which governs the vertical circulation of the oceans, known as the thermohaline circulation.


Ocean geochemistry is the discipline focusing mostly on the inorganic constituents of seawater in the world oceans. Interactions with biology and organic chemistry, and external sources and sinks, such as rivers, atmosphere, hydrothermal vents, and sediments, do play a role. Marine geochemistry comprises a far wider range including other aspects of inorganic geochemistry and organic geochemistry, and secondly not only the seawater but also investigations focusing on the underlying marine sediments and their inorganic and organic contents.

Seawater comprises above all the water molecule (H2O) and in addition all of the existing 91 chemical elements in their dissolved state in water, in concentrations ranging many orders of magnitude. Most abundant at ~0.5–0.6 moles per liter (M for molar) are the major constituents of dissolved seasalt, the ions Na+ and Cl (Table 1). These and other major elements are the topic of this chapter.
Table 1

The major ion composition of seawater at salinity S = 35.


Atomic weight (or molecular for SO4 2−)

Ocean concentration

Rivers average concentration (Molar [mol.dm−3])

Ocean residence time versus rivers input [Year]


[gram.kg−1] at S = 35

Molar [mol.L−1 = mol.dm−3]






83,900,000 (83.9 million)

SO4 2−




















8,500,000 (8.5 million)



~0.0018 (variable)







52,600,000 (52.6 million)






11,600,000 (11.6 million)






1,000,000 (1 million)






10,500,000 (10.5 million)






4,000,000 (4 million)

Latter value is close to the world average salinity 34.7 of the modern ocean. Compilation after Dittmar (1884), Wilson (1975), Quinby-Hunt and Turekian (1983), Murray (1992), and Sarmiento and Gruber (2006). Riley (1965) described the historical development of salinity and its measurement. The molar concentrations in seawater at S = 35 have been calculated from the original [gram.kg−1] divided by the atomic weight and adjusted for a “typical” density of seawater of 1027 [gram.dm−3 = kg.m−3] or 1.027 kg/Liter. This chosen density is merely an example value, in this choice for water with temperature of ~10 °C at the sea surface (density) or collected from the deep and brought up to the surface (potential density). Calcium (Ca) varies somewhat (~1%) between surface waters and deep waters. Mg and Sr are in the same group in the periodic table as Ca and also show some variations, that however are too small deviations from being “conservative” to affect salinity at its current accuracy. Boron (B) exists as both neutral B(OH)3 and negatively charged B(OH)4 ion in seawater, their relative proportion being a function of pH (Dickson 1990). At “typical” pH = ~8 for modern surface waters the concentration of B(OH)4 is insignificant for the current accuracy of measured conductivity from which salinity is calculated. Similarly, the HCO3 ion is barely significant for overall conductivity, i.e., derived salinity. For the overall world ocean, a “typical” salinity of 35 is commonly used. Similarly in laboratory experiments, and in all sorts of calculations, one commonly uses a salinity of 35

Three orders of magnitude less abundant is Dissolved Inorganic Carbon (C as DIC) varying around ~2 millimoles [mM = 10−3 M], and pivotal for life in the sea; this is treated in the chapter on “Carbonate Minerals and the CO2-Carbonic Acid System” (Lerman and Mackenzie, this volume).

Much less abundant are the nutrients nitrate and phosphate that occur in the micromole [μM = 10−6 M] range and are essential for each living organism. Also essential for life are several trace nutrient elements. The nutrients, trace nutrients, and other trace elements are treated in the chapter on “Ocean Biochemical Cycling and Trace Elements” (De Baar et al. this volume).

Most of the chemical elements have several isotopes , which can be either radioactive isotopes or stable isotopes, that provide useful insights into ocean processes and are described in other chapters.

Major Ions in Seawater and Salinity

Upon accurate analyses of 77 seawater samples, Dittmar (1884) confirmed previous findings of Marcet and of Forchhammer that, regardless of the absolute concentration of the total solids, the ratios between the more abundant major chemical elements are virtually constant (Table 1). Since then, these major constituents of seawater are classified as being “conservative.” Variation is ascribed exclusively to there being more or less pure water in anyone given sample of seawater due to net precipitation or net evaporation at the sea surface, respectively, or due to mixing of two water masses with different total salt content. These findings lead to the concept of salinity (after Forchhammer), that is the total amount of dissolved salts in seawater. Salinity is expressed as gram per kg seawater [absolute salinity; gram.kg−1] that is dimensionless and hence can also be expressed as per mille unit 0/00 or nowadays as practical salinity without any unit at all (IOC 2010).

The salinity in the world oceans varies in a narrow range between ~33 and ~38, such that 35 is taken as the “typical” value (Table 1). The nice thing of the uniform proportions is that when you measure just one, you know them all, and also their summation, the salinity. This was done for 65 years until the ~1960s by measuring the chlorine concentration and from this calculate salinity. Nowadays, the salinity is measured by the electrical conductivity of seawater that depends on the total amount of ionic charges, the more charged ions, the better the conductivity, and the higher the derived salinity. The combined and very accurate measurements by electronic sensors throughout the ~4000 m deep water column of the oceans of conductivity (for salinity); in situ temperature; and pressure (for depth) yield the three key variables Salinity, Temperature, and Pressure (S, T, p) that together determine the density [kg.dm−3] of seawater at the given location and depth. The density is calculated very accurately with the thermodynamic equation of state of seawater (IOC 2010). For the two other parameters, the temperature in the oceans varies from around −1.85 °C (the freezing point of seawater) to ~25 °C or in some tropical seas even 30 °C or higher; nevertheless, the temperature of deep ocean waters is quite uniform around ~2 °C. The pressure increases with depth to 380 atmosphere at the average ~ 3800 m depth of the oceans, or at most ~1000 atm in the deepest ~10 km deep oceanic trenches.

The variations of density are small, but extremely important as these small differences in density of different water masses are the drivers of the “thermohaline” circulation of the oceans. The density of ocean waters ranges from ~1020 to ~1030 gram.dm−3 (or kg.m−3), i.e., seawater is 2–3% more dense (heavier) than pure water. The very small range of about 1% (i.e., from 102% to 103%) of the overall density represents the complete window that drives the circulation. Therefore physical oceanographers are keen to obtain very accurate estimates of salinity, currently at around the third decimal point, e.g., 35.456 with an accuracy of ±0.001.

One must be aware of minor deviations of some chemical elements from the constant proportions, i.e., from being conservative. For example, Dittmar (1884) already noticed that the Ca2+ in deep water samples was slightly higher than in the upper waters. This is due to the net uptake from upper water by biocalcification , that is the production by organisms of solid CaCO3 for biogenic shells and corals.
Ca 2 + + 2 HCO 3 - CaCO 3 crystalline + CO 2 + H 2 O Open image in new window
Upon senescence of the shell-forming organism, the heavy (density ~ 2700 kg.m−3) CaCO3 shells settle into deep waters and beyond a given depth horizon enter waters that are undersaturated versus solid CaCO3, such that the shells slowly dissolve. The overall result is that the deep waters comprise, on average, about 1% more dissolved Ca2+ ions than the surface waters.
The relative composition of the dissolved salts in seawater appears to have been fairly very uniform over the past 541 million years to present, the Phanaerozoic Eon . Evidence for this comes from marine evaporite deposits (e.g., Ryan 2008). Occasionally a semi-enclosed sea has become completely enclosed, and during evaporation of the water, the seasalt has deposited in a well-known sequence in the reverse order of the solubilities of the minerals. Therefore the order of precipitation from sea water is:
  1. 1.

    Calcite (CaCO3) and dolomite (CaMg(CO3)2)

  2. 2.

    Gypsum (CaSO4-2H2O) and anhydrite (CaSO4).

  3. 3.

    Halite (i.e., common salt, NaCl)

  4. 4.

    Potassium and magnesium salts

Given the dominance of Na+ and Cl in seawater, the thickness of the halite layer (NaCl) amounts to some 90% of the overall deposit. When adding up the contents of all the layers, one finds the ratio of elements to be the same as in modern seawater. Given the age of the studied evaporite deposit one knows that, at that time in the past, the dissolved seasalt was the same composition as today.

For example, the Messinian Salinity Crisis (MSC) that occurred from 5.96 to 5.33 million years in the past is deemed to be caused by the closing and opening again, perhaps several times, of the Strait of Gibraltar, thus cutting off the Mediterranean Sea from the Atlantic Ocean (Hsu 1983; Roveri et al. 2014). This caused complete dessication of the Mediterranean Sea, perhaps several times, and as a result thick packages of evaporites were laid down and nowadays can be found and sampled in the sediments throughout the Mediterranean Sea. The water from the Mediterranean would have been redistributed in the world ocean, but not the salt that remained in the deposits. As a result there must have been a global decrease in salinity. The total element composition of the MSC evaporites closely confirms that of modern seawater. From worldwide estimates of all such evaporite deposits and their ages, it is concluded the seawater composition has been fairly stable during at least the past 541 million years, although some deviations are known (e.g. Blättler and Higgins 2014).

Residence Times and Mass Balance

In the hydrological cycle, the oceans lose water by evaporation and rivers return the water to the ocean. The river water contains very little dissolved salts (Table 1), and also in very different proportions than seawater. Notably in Table 1 for the positive ions in river water the predominance is
Ca 2 + > Mg 2 + > K + > Na + > Sr 2 + Open image in new window
whereas in seawater the predominance is
Na + > Mg 2 + > Ca 2 + > K + > Sr 2 + Open image in new window
For the ocean one may assume that the inflow of water does on average equal the net loss by evaporation, a stable situation and that is called a steady state . When now dividing the large ocean volume of 1330 × 106 km3 by the annual ~36 × 103 km3.year−1 freshwater inflow one obtains the ~37,000 years residence time of water in the oceans, this relative to river inflow (plus groundwater discharge).
Next one can for the major chemical elements (Table 1) multiply this water residence time with the ratio of concentrations [mol.dm−3] in seawater and in river water
Residence time = volume ocean / inflow rivers × concentration ocean / concentration rivers Open image in new window
and obtain the residence time [year] of major elements in the oceans (Table 1). These residence times are very long, from 52.6 million to 83.9 million years for major elements Na and Cl, respectively, to ~10 million years for Mg and K, but “only” 1 million year for Ca. For the latter, this results from biocalcification. While, as already mentioned, some of these shells dissolve again in the deep sea, a portion remains intact and becomes buried within marine sediments as lime deposits.
If their concentrations do not change through time, then a steady state must exist, by and large, for each of the dissolved major elements of the seasalt. Somehow the loss term from seawater must be found in marine deposits. An obvious removal term of dissolved major ions is the occasional formation of marine evaporites at various locations of the globe, such as those underlying the Mediterranean Sea. Another sink are the vast and often also quite thick calcareous minerals, the buried fossil shells of CaCO3, that also comprise some trace amounts of akin elements Mg and Sr. Also, because the crystal CaMg(CO3)2 or dolomite is thermodynamically more stable than CaCO3 (limestone ), very slow postdepositional transformation, i.e., dolomitization , occurs within the sediments.
2 CaCO 3 limestone + Mg 2 + CaMg CO 3 2 dolomite + Ca 2 + Open image in new window
As a result, the concentration of Mg2+ in pore waters of such sediments is lower than in the overlying seawater, and extra Mg2+ diffuses from the sea into the sediment pore waters. This is an extra loss term for Mg2+ from seawater.

Another sink term for Mg, and K, is their reaction with degraded clays coming from land and suspended in the river outflow, to various marine type clay minerals that in the process have taken up extra Mg and K. This “reverse weathering” hypothesis (Mackenzie and Garrels 1966) was at first difficult to test in the laboratory due to the very slow reactions. However, this was eventually demonstrated in long-term experiments on the formation of authigenic aluminosilicates in the sediments of the Amazon delta (Michalopoulos and Aller 1995). Meanwhile the discovery in 1977 and beyond of hydrothermal circulation of seawater at Mid Ocean Ridges did provide an answer at least for Mg (Corliss et al. 1979). The cold (~2 °C) deep seawater entering the rocks through cracks heats up enormously and reacts with the rocks, and the hot seawater plumes coming out at ~380 °C have lost all dissolved Mg. Hence, hydrothermalism is a major sink for removal of Mg from ocean waters.

The top 10 rivers in terms of water discharge supply the majority of all fresh water and its dissolved ions into the oceans (Martin and Meybeck 1979). For example, the Amazon, Zaire, and Mississippi rivers draining into the Atlantic Ocean, together represent almost 50% of all river inflow into the world oceans. Almost all large rivers are regularly monitored for volume flow and dissolved major ions, such that, despite interannual variability, a quite reliable worldwide estimate can be obtained, at least for our era.

Overall, the above simple approach merely comparing river supply with ocean content and loss to evaporites, limestone deposits, reverse weathering, and hydrothermalism is quite satisfactory as a steady state model for the, mostly conservative, major elements of seawater (Table 1). However, many minor and trace elements (see chapter “ Ocean Biochemical Cycling and Trace Elements”) require a more complete concept of ocean mass balance (Figure 1) including consideration of several other sources and sinks:
  • River input source

  • Aeolian supply source of terrestrial dust as either dry deposition or wet deposition

  • Gas exchange with the atmosphere notably for CO2 and O2, as source or as sink

  • Hydrothermal supply or loss, for example, an iron source (Fe), or sink of magnesium (Mg)

  • Radioactive production source or radioactive decay loss within the oceans

  • Sedimentation loss

  • Diffusive fluxes into or out of sediment pore waters

Figure 1

Box cartoon of ocean inventory box (azure color) with all sources and sinks, and in steady state ideally: Σsources  =  Σsinks, where Σsources  =  Rivers + dust + air to sea gas inflow + hydrothermal input + radioactive production + diffusion from sediment and Σsinks  =  Air to sea gas outflow + hydrothermal output + radioactive decay + sedimentation + diffusion into sediment.

This also implies that for deriving a steady state mass balance for any one chemical element of interest, one has to add up all the sources, as well as add up all the sinks. Next one somehow hopes these are equal, or adjusts these, such that the sum of sources more or less equals the sum of sinks. Finally one can compare these sums [moles.year−1] with the total ocean inventory [moles] to derive a residence time [year] for the given chemical element.

During the Last Glacial Maximum (LGM), large ice caps covered large parts of North America and Europe. The volume of these ice caps corresponded with an ~120 m lower sea level than today. As a result of the corresponding lower overall volume of the seawater in the oceans, the salinity was higher in the LGM (Adkins et al. 2002), with a different salinity distribution in the major oceans than today, but the relative proportions of dissolved salts remained the same. Modern day salinities are all within 0.5 of the global average salinity of 34.7, whereas salinities during the last glacial maximum (LGM) ranged from 35.8 in the North Atlantic to as high as 37.1 in the Southern Ocean .

Circulation of Deep Ocean Waters

The world ocean distributions of chemical elements are more strongly controlled by the circulation of deep ocean waters, that is all waters below ~500 m depth, than by the circulation of the upper ocean (all waters shallower than ~500 m). Thus in order to understand the cycling and distribution of the chemical elements, one at least also needs to have some understanding of the deep Thermohaline Circulation (Van Aken 2007).

The ultimate driver of this deep circulation is the formation of deep waters. This sinking is primarily caused by a seasonal increase of density in the winter in polar regions. Firstly by winter cooling; secondly by seasonal formation of sea-ice resulting in the remaining seawater having a higher salinity, the water overall gets a higher density.

In the Arctic region, this winter process is most intense in February in the Greenland-Iceland-Norwegian Seas (GIN-Seas; Hopkins 1991; Van Aken 2007; Seidov et al. 2013, 2014) north of Iceland. This deep basin fills up with this denser water and overflows at the relatively shallow Denmark Strait between Iceland and Greenland, and the straits between Iceland, the Faeroes, and Scotland. These overflows sink again to near the bottom of the northern North Atlantic Ocean and flow southwards, along the Western Atlantic Ocean basin due to Coriolis acceleration . This North Atlantic Deep Water (NADW) also picks up some waters that were formed in winter in the Labrador Sea (between Canada and Greenland) that are less dense and therefore sink only to ~1200 m depth. Moreover within the Mediterranean Sea there is heating and net evaporation of seawater. This warmer but more saline (S ~ 38) water sinks and eventually overflows at Gibraltar Strait into the Atlantic Ocean, where this Mediterranean Overflow Water (MOW) spreads out at ~1200 m depth as far as the West Atlantic. By vertical mixing, the MOW causes the underlying NADW to become more saline.

After some 400–500 years, the NADW enters the Antarctic Circumpolar Current . In the Antarctic Weddell Sea, there is similar formation of cold Weddell Sea Deep Water (WSDW) during the Antarctic winter (July–August). This eventually outflows into all three major ocean basins as Antarctic Bottom Water (AABW). At the Antarctic Polar Front, there is also sinking of waters but only to ~1000 m depth, this Antarctic Intermediate Water (AAIW) also outflows into all three major ocean basins. As a result of these outflows of Antarctic origin waters, the deep Indian Ocean is filled up until its boundary with Asia at 20–30 oN. The oldest deep waters of the northern Indian Ocean have an “age” of ~1000 years; in other words, it has been ~1000 years ago that this water last was at the surface. Similarly, the deep Pacific Ocean fills from south to north and in the deep North Pacific Ocean exists the oldest deep waters with ages in the 1500–1800 or even 2200 years range.

The abovementioned ages are based on the distribution of 14C in the deep oceans. More recently, Matsumoto (2007) has revised these conventional ages (his Figure 1) by accounting for two key mechanisms and has arrived at somewhat younger ages (his Figure 4), notably ~200 years for the Northwest Atlantic location, and ~1200 years respectively for the Northeast Pacific region. The age difference of ~1000 years between these two positions would then be a bit less, but still very long. Other age estimations have been made (Khatiwala et al. 2012; Gebbie and Huybers 2012) yet all agree that the age difference between deep North Atlantic and deep North Pacific Oceans is in the order of ~1000 years or more (Figure 2).
Figure 2

Simplified scheme of circulation in the deep ocean basins. Filled blue arrows are (vertical) deep water formation and (horizontal) flow. Formation of deep waters shown in red. Not shown are several other (smaller) sources, in the Labrador Sea, the Mediterranean Overflow Water (MOW), in the Ross Sea and along some other sectors around Antarctica. The open blue arrows are for general upwelling that closes the water budget by returning water to the uppermost boxes in all oceans. From the Pacific Ocean, there is a return flow (shown in purple) via the Indonesian Archipelago, Agulhas Current, Gulf Stream resuppplying the GIN-Seas. Antarctic Ocean is all waters south of the Antarctic Polar Front (APF) at ~50 oS, the APF is defined by the surface expression of the 2 °C isotherm (Deacon 1984). Drafted after similar schemes that exist in the literature (e.g. Schmitz 1996; his Fig. I-91 at page 106) and at the internet (e.g. http://ccsr.aori.u-tokyo.ac.jp/old/ehtml/eocean.html).


Salinity of seawater in the oceans comprises the small number of chemical elements that occur in very high concentration in seawater. These major elements in seawater have very long residence times in the oceans, for example, 52.6 million and 83.9 million years for major elements Na and Cl, respectively, and ~10 million years for Mg and K, but “only” 1 million year for Ca. This in comparison to the relatively very brief mixing time of seawater between the ocean basins in the order of “merely” ~1000 years yields a very constant ratio of these major elements in seawater. One exception is again Ca that due to its involvement in the biogeochemical cycle of CaCO3 shells and corals has a slightly lower concentration in surface waters relatively to deep waters. From quantitative studies of worldwide evaporite deposits, it has been estimated that the seawater composition has been fairly stable during at least the past 541 million years. In the modern ocean, the “thermohaline” circulation of deep waters is driven by small differences of seawater density due to differences in temperature and salinity.

Winter cooling and sea-ice formation in polar winters cause surface waters to become more dense and sink into the deep polar seas, from where these waters outflow into all other oceans.



  1. Adkins JF, McIntyre K, Schrag DP (2002) The salinity, temperature, and delta 18O of the glacial deep ocean. Science 298(5599):1769–1773CrossRefGoogle Scholar
  2. Blättler CL, Higgins JA (2014) Calcium isotopes in evaporites record variations in Phanerozoic seawater SO4 and Ca. Geology 42:711–714CrossRefGoogle Scholar
  3. Corliss JB, Dymond J, Gordon LI, Edmond JM, Von Herzen RP, Ballard RD, Green K, Williams D, Bainbridge A, Crane K, Van Andel TH (1979) Submarine thermal springs on the Galapagos Rift. Science 203:1073–1083CrossRefGoogle Scholar
  4. Deacon GER (1984) The Antarctic circumpolar ocean. Cambridge University Press, Cambridge, 180pGoogle Scholar
  5. Dickson AG (1990) Thermodynamics of the dissociation of boric acid in synthetic seawater from 273.15 to 318.5 K. Deep-Sea Res 37:755–766CrossRefGoogle Scholar
  6. Dittmar W (1884) Report on researches into the composition of ocean water, collected by H.M.S. Challenger. Chall Rep Phys Chem 1:1–251Google Scholar
  7. Gebbie G, Huybers P (2012) The mean age of ocean waters inferred from radiocarbon observations: sensitivity to surface sources and accounting for mixing histories. J Phys Oceanogr 42:291–305CrossRefGoogle Scholar
  8. Hopkins TS (1991) The GIN sea – a synthesis of its physical oceanography and literature review 1972–1985. Earth Sci Rev 30(3–4):175–318CrossRefGoogle Scholar
  9. Hsu KJ (1983) “A voyage of the Glomar challenger”. The Mediterranean was a desert. Princeton University Press, PrincetonGoogle Scholar
  10. IOC (2010) The international thermodynamic equation of seawater – 2010: calculation and use of thermodynamic properties. Intergovernmental oceanographic commission, manuals and guides no. 56, UNESCO (English), 196 p. Free download at: http://www.teos-10.org/pubs/TEOS-10_Manual.pdf
  11. Khatiwala S, Primeau F, Holzer M (2012) Ventilation of the deep ocean constrained with tracer observations and implications for radiocarbon estimates of ideal mean age. Earth Planet Sci Lett 325–326:116–125CrossRefGoogle Scholar
  12. Mackenzie FT, Garrels RM (1966) Chemical mass balance between rivers and oceans. Am J Sci 264:507–525CrossRefGoogle Scholar
  13. Martin JM, Meybeck M (1979) Elemental mass-balance of material carried by major world rivers. Mar Chem 7:173–206CrossRefGoogle Scholar
  14. Matsumoto K (2007) Radiocarbon-based circulation age of the world oceans. J Geophys Res 112:C09004.  https://doi.org/10.1029/2007JC004095 Google Scholar
  15. Michalopoulos M, Aller RC (1995) Rapid clay mineral formation in Amazon delta sediments: reverse weathering and oceanic elemental cycles. Science 270:614–617CrossRefGoogle Scholar
  16. Murray JW (1992) The oceans. In: Butcher SS, Charlson RJ, Orians GH, Wolfe GU (eds) Global biogeochemical cycles. Academic, London, pp 175–211CrossRefGoogle Scholar
  17. Quinby-Hunt MS, Turekian KK (1983) Distribution of elements in sea water. Eos Trans AGU 64:130–131CrossRefGoogle Scholar
  18. Riley JP (1965) Historical Introduction. Chapter 1 in: Riley JP, Skirrow G (ed) Chemical Oceanography, Academic Press, London and New York, vol. 1st edn. pp 1–41.Google Scholar
  19. Roveri M, Flecker R, Krijgsman W, Lofi J, Lugli S, Manzi V, Sierro FJ, Bertini A, Camerlenghi A, De Lange G, Govers R, Hilgen FJ, Hübscher C, Meijer PT, Stoica M (2014) The Messinian salinity crisis: past and future of a great challenge for marine sciences. Mar Geol 352:25–58CrossRefGoogle Scholar
  20. Ryan WBF (2008) Modeling the magnitude and timing of evaporative drawdown during the Messinian salinity crisis. Stratigraphy 5(3–4):227–243Google Scholar
  21. Sarmiento JL, Gruber N (2006) Ocean biogeochemical dynamics. Princeton University Press, Princeton, xii + 503pGoogle Scholar
  22. Schmitz, WJ Jr (1996) On the world ocean circulation, vol. I. Technical Report WHOI-96-03; page 106 of 148pGoogle Scholar
  23. Seidov D, Baranova OK, Biddle M, Boyer TP, Johnson DR, Mishonov AV, Paver C, Zweng MM (2013) Greenland-Iceland-Norwegian seas regional climatology. Regional Climatology Team, NOAA/NODC.  https://doi.org/10.7289/V5GT5K30. See also: https://www.nodc.noaa.gov/OC5/regional_climate/gin-seas-climate/
  24. Seidov D, Antonov JI, Arzayus KM, Baranova OK, Biddle M, Boyer TP, Johnson DR, Mishonov AV, Paver C, Zweng MM (2014) Oceanography north of 600N from world ocean database. Prog Oceanogr.  https://doi.org/10.1016/j.pocean.2014.02.003
  25. Van Aken H (2007) The oceanic thermohaline circulation: an introduction. Springer, New York, XVIII + 326pGoogle Scholar
  26. Wilson TRS (1975) Salinity and the major elements of seawater. Chapter 6 In: Riley JP, Skirrow G (ed) Chemical oceanography, Academic Press, London, New York, San Francisco, vol. 1, 2nd edn. pp 365–413Google Scholar

Copyright information

© Springer International Publishing AG 2017

Authors and Affiliations

  • Hein J. W. de Baar
    • 1
    Email author
  • Steven M. A. C. van Heuven
    • 1
  • Rob Middag
    • 1
  1. 1.NIOZ Royal Netherlands Institute for Sea Research, Department of Ocean Systems (OCS), and Utrecht UniversityDen BurgThe Netherlands