Encyclopedia of Geochemistry

Living Edition
| Editors: William M. White

Ocean Biochemical Cycling and Trace Elements

  • Hein J. W. de Baar
  • Steven M. A. C. van Heuven
  • Rob MiddagEmail author
Living reference work entry
DOI: https://doi.org/10.1007/978-3-319-39193-9_356-1

Definition

Ocean biochemical cycling refers to the distribution of nutrients and bio-essential elements at low concentration that is controlled by their uptake by phytoplankton in surface waters, sinking and remineralization of organic remains in deeper waters, and subsequent redistribution by thermohaline circulation.

Introduction

Dissolved inorganic carbon (C as DIC) varies around ~2 millimoles [mM = 10−3 M] in seawater and is pivotal for life in the sea. 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 , notably Fe and Zn, that occur in the nanomolar [nM = 10−9 M] range or even lower such as Co in the picomolar [pM = 10−12 M] range. Several other trace elements also occur in the nanomolar [nM] to picomolar [pM] range. Finally a few ultratrace elements occur in the femtomolar [fM = 10−15 M] range in seawater.

The distribution of many of the chemical elements in seawater is strongly controlled by biological processes, not only in the modern ocean but also due to the major role of biological evolution in the chemical evolution of the oceans (Vernadsky 1926; Redfield 1958; Lovelock 1979; Holland 1984).

Ocean Biological Cycle Requires Bio-essential Elements

The primordial atmosphere and ocean did not comprise free O2 (see Precambrian Geochemistry ). In the mildly reducing Archean ocean (~3.8–2.5 Gy before present), the transition metals manganese (Mn), iron (Fe), cobalt (Co), nickel (Ni), copper (Cu), and zinc (Zn) would have been readily soluble as Mn2+, Fe2+, Co2+, Ni2+, CuCl2−, and Zn2+, respectively (De Baar and La Roche 2003; Saager et al. 1989; Lewis and Landing 1991), and would no doubt have been present at relatively high concentrations in this primordial ocean where reduced sulfide was also present (Turner et al. 2001). As the first life evolved, these transition metals were incorporated in many biochemical functions. As a result, all modern cellular organisms living today, prokaryotes and eukaryotes, comprise and hence require the following chemical elements in order of their typical abundance in the living cell:
$$ \mathrm{C}>\mathrm{N}>\mathrm{P}>>>\mathrm{Fe}>\mathrm{Zn}>\mathrm{Mn}>\mathrm{Cu}>\mathrm{N}\mathrm{i}>>\mathrm{Co} $$
(1)
The major biochemical functions are listed in Table 1.
Table 1

The nine chemical elements that are essential for life and their major biochemical functions, their symbols, and their names in bold print. Essential also relates to the concept that each one of these can occasionally be a limiting element for growth in conditions that may occur of very low availability, i.e., low abundance as per the concepts of limitation after Von Liebig (De Baar 1994). Please be informed that there exist several different types of superoxide dismutases, all with the same function but with quite different metal elements as cofactor. Not listed are hydrogen (H) that occurs in every biological molecule and oxygen (O) that also is very common notably in various organic functional groups; both are so very abundant in the water molecule (H2O), i.e., never potentially limiting, that they are not being treated as essential elements. Sulfur is listed here, but is abundant as major sulfate ion in seawater and never bio-limiting. The uptake of Mg for chlorophyll is trivial compared to the very high background concentration at ~0.48 mol.L−1 of Mg in ambient seawater (Reynolds 2006). Next to the nine elements essential for all life (bold print) and S and Mg, also listed are six more elements that play a role in specific taxonomic groups, among which Si and Ca are most important for formation of hard parts (e.g., shells, frustules, corals) and Mo for N2 fixation by diazotrophs. Compilation after Twining and Baines (2013) also based on Sunda (1989), Michibata and Sakurai (1990), Frausto da Silva and Williams (1994, 2001), De Baar and La Roche (2003), Wolfe-Simon et al. (2005, 2006), Bruland and Lohan (2003), Bruland et al. (2014)

Atomic number

Mass

Symbol

Name

Major biochemical function

6

12.011

C

Carbon

Carbohydrates and in every biochemical molecule

7

14.007

N

Nitrogen

Amino acids in proteins and many other molecules and functions

12

24.305

Mg

Magnesium

Central atom in chlorin ring of chlorophyll molecule, active site of RuBisCO

14

28.086

Si

Silicon

Opaline (SiO2) external frustules of diatoms, skeletons of radiolaria

15

30.974

P

Phosphorus

Phospholipids, ATP

16

32.065

S

Sulfur

Methionine, cysteine, ferredoxin

20

40.078

Ca

Calcium

Calcite and aragonite skeletons and shells, coral reefs

23

50.942

V

Vanadium

Unique high concentration in blood cells of Ascidiacea (sea squirts) (Michibata and Sakurai 1990)

25

54.938

Mn

Manganese

O2-evolving enzyme, superoxide dismutase, arginase, phosphotransferases

26

55.854

Fe

Iron

Electron transport in photosynthesis and respiration, nitrate and nitrite reductases, N2 fixation: conversion of hydrogen peroxide to water, peroxidase reduction of reactive oxygen species, superoxide dismutase

27

58.933

Co

Cobalt

Vitamin B12 in C and H transfer reactions

28

58.693

Ni

Nickel

Urease hydrolysis of urea, superoxide dismutase

29

63.546

Cu

Copper

Photosynthesis electron transport, mitochondrial electron transport, ascorbic acid redox, superoxide dismutase

30

65.39

Zn

Zinc

Carbonic anhydrase, alkaline phosphatase, RNA polymerase, tRAN synthetase, reverse transcriptase, carboxypeptidase, superoxide dismutase

42

95.94

Mo

Molybdenum

N2 fixation, nitrate reductase

48

112.41

Cd

Cadmium

Carbonic anhydrase

The initially reducing ocean environment eventually became oxidizing as more and more free O2 was generated by photosynthesis (Crowe et al. 2013) and the bio-essential trace elements, notably Fe and Mn, precipitated by oxidation. Notably Fe, which once was present abundantly in seawater and therefore incorporated in many biochemical functions during evolution, nowadays is in very low concentrations in seawater and often the limiting bio-essential element for growth of phytoplankton. Thus by producing O2, the algae eventually themselves caused their own Fe limitation.

Within the marine ecosystem of today, the unicellular phytoplankton or algae and the zooplankton and bacterioplankton are the most abundant in terms of biomass. The net growth or net decomposition of the overall plankton pool corresponds to net loss or gain, respectively, of significant amounts of these nine bio-essential elements in the ambient seawater. The more the biology has an impact on the oceanic presence of an element, the more the distribution of such element deviates from salinity . Therefore the ocean distributions of these nine bio-essential elements are nonconservative, i.e., deviating from the constant proportions of the major elements that constitute the overall salinity (see entry Ocean Salinity, Major Elements, and Thermohaline Circulation).

Other Chemical Elements with a Biological Role

In addition several taxa of marine organisms also incorporate other chemical elements for biological functions (Table 1), and this may, more or less, affect the ocean distribution of such element.

Sulfur (S) is utilized by most organisms, but the background concentration of sulfate in seawater is very high and not really affected. Similarly the uptake of Mg for plant chlorophyll hardly affects the very high background concentration of Mg in seawater.

Bio-calcification is by organisms that secrete hard shells and skeletons of CaCO3 in either crystalline forms of either calcite or aragonite. By removal of both Ca2+ and HCO3 from seawater, both are also somewhat nonconservative, somewhat because this biological role is small relative to the high abundance of Ca2+ and to a lesser extent HCO3 in seawater. The major calcite producers in the open ocean are the group of coccolithophorid phytoplankton and the foraminifera zooplankton . The most abundant pelagic aragonite producers are the pteropods , and aragonite also forms naturally in corals and in almost all mollusk shells.

The other major type of hard parts is the opaline silicon oxide (SiO2) that is produced by the important group of the diatoms among the phytoplankton, which are deemed to represent in the order of ~40% of all primary production in the oceans, and the radiolaria among the zooplankton. The concentration of dissolved silicate in seawater is relatively modest in the micromole [μM = 10−6 M] range and nonconservative due to this role in biology. Therefore silicate is commonly classified with the other major nutrients nitrate and phosphate.

There are three more trace elements that have a known biological function in some organisms. Vanadium (V) has very rare biological functions in ascidians (sea squirts) and a few other organisms (Michibata and Sakurai 1990). Molybdenum (Mo) is essential for nitrogen fixation that is one of the important biochemical pathways that play a role in controlling the oceanic nitrogen inventory (La Roche and Breitbarth 2005). Cadmium (Cd) has been found to serve as the pivotal atom in the enzyme carbonic anhydrase of a few diatom species investigated thus far (Xu et al. 2008).

Brief Outline of Marine Ecosystem Cycling of Elements

The basis of the marine ecosystem is photosynthesis, i.e., the primary production in the oceans. This is by either (i) single-cell phytoplankton or (ii) multicellular macroalgae or (iii) the symbiontic zooxanthellae within coral structures. The single-cell phytoplankton can be either prokaryotes , that is, cells without a nucleus, or eukaryote cells that do have a nucleus. Multicellular macroalgae and single-cell zooxanthellae both are eukaryotes. Marine photosynthesis in the oceans tends to follow the overall reaction
$$ {\displaystyle \begin{array}{l}106\mathrm{DIC}+122{\mathrm{H}}_2\mathrm{O}+16{{\mathrm{NO}}_3}^{-}+1\mathrm{H}{{\mathrm{PO}}_4}^{2-}\hfill \\ {}+\left(\mathrm{Mn},\mathrm{Fe},\mathrm{Co},\mathrm{Ni},\mathrm{Cu},\mathrm{Zn}\right)\hfill \\ {}+\mathrm{energy} \rightleftarrows {\left[{\left({\mathrm{CH}}_2\mathrm{O}\right)}_{106}{\left({\mathrm{NH}}_3\right)}_{16}{\left(\mathrm{H}{\mathrm{PO}}_4\right)}_1\left(\mathrm{Mn},\mathrm{Fe},\mathrm{Co},\mathrm{Ni},\mathrm{Cu},\mathrm{Zn}\right)\right]}_{\mathrm{organic}\ \mathrm{biomass}}\\ {}+ 138{\mathrm{O}}_2\hfill \end{array}} $$
(2)

adapted after Redfield, Ketchum, and Richards (1963) but here also including the six bio-essential trace elements (Mn, Fe, Co, Ni, Cu, Zn) for which their very low relative amount is subject of research nowadays (see below section Trace Nutrient Elements). The resulting organic biomass is energetically rich and hence serves as a food and energy source for bacteria and animals in the reverse reaction (2), i.e., remineralization or respiration .

In a real phytoplankton bloom in surface waters or in the ensuing export of biogenic debris into the deeper waters, the element ratio values may deviate considerably from the values given in above Eq. 2. However due to deep ocean mixing, such deviations tend to “average” such that at any given station, the proportions are more close to those in Eq. 2, but still significant variations exist.

The above stoichiometric constants (i.e.,106, 16, 1, −138), known as the Redfield ratios, were at the time largely based on concentrations measured mostly in the North Atlantic region. Since then variations in ocean distributions of, for example, nitrate and phosphate, and the causes thereof have been discussed on the basis of then available larger more worldwide datasets (Fanning 1992; De Baar et al. 1997), and significant adjusted values of the stoichiometric constants have been derived and advocated (e.g., Anderson and Sarmiento 1994), as reviewed by Sarmiento and Gruber (2006) and others.

Photosynthesis occurs only within the upper euphotic zone of the oceans where it produces major amounts of dissolved oxygen O2, which can become a few percent oversaturated in daytime during an intense phytoplankton bloom. The euphotic zone is defined as the depth the zone where sunlight is available to the 1% light depth as its lower boundary. In very clear waters, the depth of the euphotic zone can be as deep as 70–80 m. However due to the abundance of algae and other plankton, this is often much less due to “self-shading,” while in nearshore waters, the presence of mineral particles (sand, silt) restricts light penetration to just a few meters or less.

The reverse reaction is respiration by either bacteria (here including the Archaea) or zooplankton and higher animals up the food chain to fish, octopus, sharks, and whales. This takes place throughout the water column. Most respiration takes place within the upper water column (<100 m depth) and remineralizes in the order of ~90% of all photosynthetic produced biomass. This respiration consumes dissolved O2, but this is rapidly replenished from the overlying atmosphere. Also C, N, and P and Fe, Zn, Cu, Mn, Ni, and Co are returned to the seawater and can be used once again for photosynthesis. Overall this is a quite efficient ~90% recycling of resources.

However some ~10% of biomass settles, as mostly dead debris, into the intermediate waters (100–1000 m depth range) and here is remineralized almost completely. Of this remaining 10%, around ~9% of total primary production is remineralized, dissolved O2 is consumed, and the concentration of O2 decreases significantly. On the other hand, DIC, nitrate, phosphate, Fe, Zn, Cu, Mn, Ni, and Co are replenished again. Finally the ~1% of primary production that escapes remineralization within the upper 1000 m settles into the deep ocean and is the food supply for all life in the deep waters and in the sediments. Life in deep waters consumes some 0.9% and leaves merely ~0.1% for the benthic biota at the seafloor that consumes ~0.09% such that merely the leftover ~0.01% is buried more or less forever within the sediments.

When now combining (Fig. 1) this biological cycling of elements with the thermohaline circulation of the deep ocean, it is realized that throughout the long journey (lasting 1000–2000 years) of the deep water from the northern North Atlantic, around Antarctica, and eventually ending up in the North Pacific, there is a steady raindown of biogenic debris settling from the surface ocean into the intermediate and deep waters.
Fig. 1

Combination of ocean circulation (blue arrows) with vertical settling of biogenic debris (green arrows) to explain the concept of accumulation of remineralized DIC and nutrients when deep water becomes older. This figure is redrafted in color after original of Broecker and Peng (1982). Formation of deep water in the Greenland-Iceland-Norwegian (GIN) Seas, see entry on Ocean Salinity, Major Elements, and Thermohaline Circulation. Notice that for simplicity of the graph, the deep water formation in the Antarctic Ocean is not shown

Within the intermediate and deep waters, the debris is remineralized to dissolved DIC and nutrients, and O2 is consumed (Eq. 2 in reverse direction). The net result is a steady accumulation of DIC and nutrients, and steady decrease of dissolved O2, from the relatively “young” northern North Atlantic Deep Water to eventually the quite “old” North Pacific Deep Water (Fig. 2).
Fig. 2

Vertical distributions of the DIC, dissolved oxygen, and the major nutrients nitrate, phosphate, and silicate in the Northwest Atlantic Ocean (blue) versus the Northeast Pacific Ocean (red). Higher DIC in deep Pacific versus deep Atlantic is due to (i) respiration of organic matter plus (ii) dissolution of CaCO3. Higher nitrate and phosphate in deep Pacific versus Atlantic due to respiration, also explaining the opposite lower dissolved oxygen. Data for Bermuda Atlantic Time-Series Station (BATS) of GEOTRACES cruise GA02-64PE321 aboard RV Pelagia, station 21 (31°45.92′N, 64°04.95′W at 13 June 2010), after Rijkenberg et al. (2015), is available at (www.geotraces.org). Data for North Pacific is from RV Melville cruise 318M2004 along WOCE line PO2, station 119 (30.00°N,159.70°W at 4 August 2004), is available in GLODAP-2 via CCHDO. (https://cchdo.ucsd.edu/cruise/318M200406). Circulation age (Matsumoto 2007) of the deep waters (>1500 m) at the North Atlantic station is ~200 years; the 14C age is ~800 years. Circulation age of the deep waters (>1500 m) at the North Pacific station is ~1100 years; the 14C age is ~2100 years

Major Nutrient Elements: Nitrate, Phosphate, and Silicate

Due to net photosynthesis (Eq. 2) in surface waters, the dissolved nitrate (NO3 ) and phosphate (HPO4 2−) are removed, and very low, often bio-limiting, concentrations remain in surface waters of the temperate and tropical oceans (Fig. 2). In contrast these major nutrients are not fully depleted in surface waters of firstly the Antarctic Ocean and also the equatorial Pacific and the Subarctic North Pacific. These ocean regions represent some 30–40% of the world oceans and are known as the high-nutrient low-chlorophyll (HNLC) waters. For these HNLC waters, we now know that the combination of iron limitation and light limitation prevents the regional phytoplankton to develop blooms to full depletion of the major nutrients. For this HNLC condition, see the below section on “Trace Nutrient Iron (Fe).”

As noted earlier, silicate is an important nutrient for diatoms, which undergo seasonal blooms. At the end of winter, there are ample N, P, and Si in surface waters, and once more sunlight appears in early spring, massive diatom blooms take place. Rapidly within 3–4 weeks, all dissolved Si is removed from surface waters, and the spring diatom blooms collapse due to Si limitation.

Another aspect is that the opaline frustule remains intact in seawater with a low dissolved silicate concentration that actually is strongly undersaturated versus opal. This is because the living diatoms have an organic coating on the frustules, preventing direct contact with the undersaturated seawater and hence preventing dissolution. Yet upon senescence the diatom debris will settle downward and hence export into deeper waters. Meanwhile the organic coating also disappears, e.g., due to bacterial consumption, and the frustules tend to dissolve. The rate of dissolution depends strongly on the surface/volume ratio such that the smaller frustules of small diatoms rapidly disappear due to dissolution, but very large diatom frustules may rapidly (~100 m/day or more) sink all the way down the ~4 km deep water column to the seafloor. Thus overall the dissolution of settling diatom frustules is a balancing act between dissolution and settling, where size plays a key role. As a result the apparent typical depth of opal dissolution is deeper than for respiration input of nitrate and phosphate, and overall the concentration of dissolved silicate increases more steadily with depth (Fig. 2). Again in combination (Fig. 1) with deep circulation, the dissolved silicate in the deep North Pacific is about threefold higher than in the deep North Atlantic (Fig. 2).

At locations where intense blooms of large-size diatoms occur in surface waters, the rain rate of settling opal dominates the sedimentation at the seafloor. Here thick layers of opaline sediments accumulate. Even bottom waters with the highest dissolved silicate at ~150 μM still are undersaturated versus opal, such that the deposit tends to dissolve. This leads to even higher concentrations in the pore waters within the sediments, and the pore water reaches equilibrium such that from there on, the opal remains preserved.

Within the Antarctic Ocean, there is major upwelling of deep waters, bringing ample supply of major nutrients N, P, and Si to the euphotic zone in the surface waters. The circumpolar Antarctic polar Front is the major region where large diatoms utilize this ample supply and grow intensely (Smetacek et al. 1997), and as a result, the underlying sediments comprise thick layers of opal, representing the major (~70%) ocean sink of dissolved Si. When comparing the river input of dissolved Si with the overall dissolved Si budget of seawater, one derives an oceanic residence time of ~10,000 years for Si (Tréguer and Rocha 2013), as compared to the ~1000 years’ time of ocean mixing. This implies that, on average, once a Si atom has entered the oceans via rivers, it will be at the surface once every ~1000 years and from there on take part in the cycle of diatom growth and dissolution. This cycle will occur some nine to ten times before the SiO2 is buried “forever” in the thick layers of opaline diatom ooze.

It has been estimated that worldwide, the cycles of N and Si are of comparable magnitude, i.e., their average ocean cycling is in a Si:N = 1:1 ratio or thereabout (Sarmiento and Gruber 2006). However it is less well realized that large variations exist; most notably the Antarctic diatoms tend to be very heavily silicified such that a ratio Si:N in the 2–3 range is more valid here. Given the key role of large Antarctic diatoms in the global Si cycle (Tréguer and Rocha 2013), and the key role of the Antarctic Ocean in world circulation, this should not be, but often is, overlooked in simulation modeling of the world ocean biogeochemical cycles.

Trace Nutrient Elements

Each of the trace elements in the oceans does have one or more sources and one or more sinks, respectively, briefly summarized in Panel A (also see entry Ocean Salinity,Major Elements, and Thermohaline Circulation, its Fig. 1).

The importance in biology of the six bio-essential trace metal elements (Table 1) is in the order Fe > Zn > Mn > Cu > Ni >> Co (as per Eq. 1). However here Fe and Mn are together discussed first, given the great significance of redox chemistry and hydrothermal supply for both elements. Next Zn, Cu, Ni, and Co are discussed and finally the three minor “biological” trace metal elements V, Mo, and Cd that play a special role for some, but by no means all, organisms.

Iron (Fe) is by far the most important trace metal transition element in biology, with many biochemical functions. Here the relative ease of redox between Fe(II) and Fe(III) is pivotal in, for example, electron transfer within the living cell (Table 1).

In the modern O2-rich ocean waters, the dissolved Fe in 0.2 μm filtered seawater is in the Fe(III) oxidation state and is predominantly (~99%) bound to organic ligands forming Fe(III)-organic complexes. The remaining ~1% is dominated by the 0.92% inorganic species Fe(OH)3(dissolved) and further 0.045% Fe(OH)4 , 0.037% Fe(OH)2 +, and merely 3 × 10−11% free Fe3+ ion. In the same O2-rich seawater conditions, the thermodynamics of the redox half reaction
$$ {\mathrm{Fe}}^{2+} \rightleftarrows {\mathrm{Fe}}^{3+}+e $$
(3)

predicts the ratio Fe2+/Fe3+ of these free ions is 100.42 or 2.63 (De Baar and De Jong 2001). Taking this ratio and given the above speciation of Fe(III) states and the speciation of Fe(II) states (Panel B), the overall ratio Fe(II)/Fe(III) is in the order of 10−12 in O2-rich seawater. When now taking a mildly anoxic environment without O2 and with a sulfide concentration of 10 μM such as in pore waters of reduced marine sediments, Eq. 3 now yields the free ion ratio Fe2+/Fe3+ of 1017.4. Taking into account the speciation, the overall ratio Fe(II)/Fe(III) is in the order of 107 in mildly reducing anoxic pore waters. Thus from open ocean waters to mildly anoxic pore waters, the ratio is shifted by a dramatic factor of 1019. This major shift is the key to understanding the driving redox forces for the cycling of Fe in the oceans. Finally, deeper in the sediments at higher S2− concentrations, there occurs precipitation of various iron sulfide deposits, thus removing Fe from solution in pore waters. Thus there is an optimal depth within the sediment where O2 is absent and S2− is still quite low where mobilization of Fe is optimal for Fe diffusing upward into overlying waters.

Panel A. External sources and sinks of dissolved trace metals in ocean waters

The external sources and sinks, and the several processes within the ocean waters (Panel B), all may affect the distribution of any given trace element in ocean waters.

River inflow is modified by estuarine mixing of the fresh river water with saline seawater. As a result there is massive flocculation of colloids (defined in Panel B) from the original river water. Most of the river inflow of many trace metals, notably Fe, is removed and deposited in estuarine sediments, such that only a small percentage truly enters the oceans but still can be traced over long distances (Klunder et al. 2012b).

Aeolian deposition into surface waters of dust coming from the continents may be either as wet deposition with rainwater or as dry deposition. Rainwater has a wide range of pH values with final deposition in the pH range 4–7, i.e., acidic and more favorable for partial dissolution of metals from dust than surface seawater at pH ~8 (Jickells and Spokes 2001).

Oxidation and reduction (redox) affects those metals that have two or more oxidation states within the overall redox range of the oceans, sediments, and hydrothermal vents. In the modern O2-rich ocean, the major dissolved states Mn2+ and Fe3+ are thermodynamically unstable and tend to deposit solid oxyhydroxides such as MnO2 and Fe(OH)3 but more often some poorly defined amorphous Fe-Mn oxyhydroxide phases. Within marine sediments the respiration of organic matter (reverse Eq. 2) by bacteria consumes the available O2 in pore waters, and supply from overlying waters is slow due to restricted vertical diffusion. Once the dissolved O2 is almost all removed, other specialized bacteria take over and use other oxidants (electron acceptors) that are more favorable under the conditions. These are in sequence first the dissolved nitrate NO3 being reduced to ammonia (NH4OH) with nitrite (NO2 ), nitrous oxide (N2O), and N2 as intermediates or also end products. Once nitrate is depleted, there secondly follows the reduction of solid MnO2 to high concentrations of dissolved Mn2+ and thirdly reduction of solid Fe(OH)3 to high Fe2+ concentrations. These very high concentrations of Mn2+ and Fe2+ in the pore waters diffuse upward in bottom waters from where they are by ocean lateral mixing transported to the open ocean, with steady loss due to the very slow or slow oxidation of Mn2+ and Fe2+, respectively. Similar redox chemistry is known to take place for dissolved cerium Ce3+ versus solid oxidized CeO2. Deeper in the sediment, there is reduction of the abundant sulfate SO4 2− to the reduced sulfite S2− state. Latter sulfide leads to precipitation of generally very insoluble FeS, CuS, ZnS, and CdS.

Hydrothermal vent effluents at mid-ocean ridges have a high temperature in the 300–400 °C range and are very strongly reducing. This is accompanied by very high concentrations of reduced Mn2+ and Fe2+ as well as elevated concentrations of dissolved Co, Cu, Zn, Ag, Cd, and Pb (Von Damm et al. 1985). Once in contact with ambient, cold (2 °C), O2-rich seawater, most of the Fe and Mn precipitate again as oxyhydroxides, also including the other metals, this leading to metalliferous sediments on the crest of mid-ocean ridges.

For example, near Elephant Island, Antarctica, there is a distinct plume of dissolved Fe coming off the extensive shelf sediments, clearly discernible and accompanied by higher levels of Mn and Al versus the low background concentrations (Fig. 3).
Fig. 3

Distinct plumes of coinciding elevated dissolved Fe, Mn, and Al coming off the shelf near Elephant Island, Antarctica. Concentrations in nanomolar per liter [nM = 10−9 Mol.L−1]. Vertical axis is sigma for density (e.g., sigma = 27.6 for density = 1027.6 kg.m−3) increasing steadily with increasing depth. The maximum concentrations are at ~96 m depth (After Klunder et al. (2014). Data in Geotraces IDP 2014 (www.geotraces.org))

When the bottom waters overlying anoxic sediments also have a low O2 content, such as in the 100–1000 m depth range of the Northwest Indian Ocean or in the subsurface waters at the East Equatorial Pacific, this dissolved Fe can be maintained at elevated concentrations in the order of 1–6 nM (Landing and Bruland 1987; Saager et al. 1989; Vu and Sohrin 2013).

In the high temperatures (300–400 °C) of extremely reducing hydrothermal effluents, the concentration of dissolved Fe may be as high as 1–2.5 millimolar (e.g., Von Damm et al. 1985, 1998), i.e., some three to eight million times higher than ambient deep ocean water. However due to mixing with ambient cold and oxidizing seawater, this is rapidly removed by oxidation and sedimentation on the ridge crest forming metalliferous sediments. Nevertheless the hydrothermal Fe source is so highly enriched that, despite the major loss at the ridge crest sediments, the plume can be traced at millionfold lower 1–3 nM Fe concentrations over hundreds of kilometers versus the low background concentration of Fe at ~0.3–0.4 nM in deep waters. For example, at the Antarctic Zero Meridian, there is an extensive ~1200 km wide plume of hydrothermal Fe near the Bouvet triple junction where three mid-ocean ridges meet (Fig. 4) and where undoubtedly intense hydrothermal effluents must exist. This plume is accompanied with high Mn also from hydrothermal input. Similarly at the Arctic Gakkel Ridge, an extensive hydrothermal plume of both Fe and Mn was found (not shown; Klunder et al. 2012a, Middag et al. 2011a), extending more than 500 km. Similarly extended plumes have been recently reported for the Pacific Ocean (Resing et al. 2015). Over that extent the concentrations of dissolved Fe and Mn steadily diminish due to combination of continued oxidation as well as dilution by mixing with ambient waters low in Fe and Mn. The decrease of dissolved Fe continues until a threshold concentration is reached that equals the concentration of Fe-binding organic ligands that from there on maintain the remaining dissolved Fe in solution. For assessing the net hydrothermal supply of Fe to the world oceans at large, this far-field extent of the relatively low-level (1–3 nM) plumes (Klunder et al. 2011, 2012a, 2014) with their slow chemical changes is most relevant.
Fig. 4

Upper graph: dissolved Fe at a 3000 km long section at the Antarctic Zero Meridian after Klunder et al. (2011). Vertical axis is pressure [decibar] that is virtually identical to depth [meter]. Notice very low Fe in surface waters notably near the Antarctic continent, vertically separated from the enriched hydrothermal plume at great depth >1000 m. Lower graph: dissolved Mn along the same section after Middag et al. (2011b). Notice very low Mn in surface waters notably near the Antarctic continent, vertically separated from the enriched hydrothermal plume at great depth >1000 m. The elevated Mn at the surface at 54 °S is due to a recent aeolian dust event, also found for Fe at the surface (less distinct in the above contour plot) and dissolved Al at the surface (not shown) (Data in Geotraces IDP 2014 (www.geotraces.org). Graphics www.eGeotraces.org using Ocean Data View (R. Schlitzer))

The longest ocean trace element section thus far is in the West Atlantic Ocean and nicely shows all the processes contributing to the distribution of dissolved Fe (Fig. 5). There is a hydrothermal plume at ~2500 m depth just south of the equator that despite the long westward distance away from the Mid-Atlantic Ridge still is discernible versus the low 0.4–0.6 nM background deep dissolved Fe. Just north of the equator (~10 °N), the subsurface Fe maximum at 200–1000 m depth coincides with a strong O2 minimum suggesting a reductive source of the elevated Fe. These O2 minimum waters flow westward (“into the paper”), then turn around clockwise, and more northerly (~40 °N) become an eastward return flow that still carries the elevated Fe signal.
Fig. 5

Section of dissolved Fe in the West Atlantic Ocean, Geotraces section GA02 (After Rijkenberg et al. 2014. Graphics using Ocean Data View (R. Schlitzer). All data as shown above by Rijkenberg and coworkers is in Geotraces IDP 2014 (www.geotraces.org))

Sahara dust input causes elevated dissolved Fe in surface waters in the 20–30 °N region, accompanied by higher Mn and Al as well, where Al is the classical tracer of dust supply (Fig. 6). This high Fe in otherwise oligotrophic (nutrient-depleted) surface waters does support N2 fixation by Trichodesmium phytoplankton in the Sargasso Sea near Bermuda.
Fig. 6

Surface water distributions of dissolved Fe, Al, and Mn along the West Atlantic section Geotraces GA02 (After Rijkenberg et al. 2014. Data is in Geotraces IDP 2014 (www.geotraces.org))

Iron Limitation of High-Nutrient Low-Chlorophyll (HNLC) Regions

There are three regions where major nutrients N, P, and Si are in ample supply in surface waters, but still phytoplankton blooms are not able to utilize this resource. These are the Southern Ocean and the equatorial Pacific and Subarctic North Pacific, together representing ~30% of ocean surface waters. The hypothesis of Fe limitation of these HNLC regions was first tested in Fe addition experiments in bottle incubations in the Subarctic North Pacific (Martin and Fitzwater 1988) and in the Southern Ocean (De Baar et al. 1990; Buma et al. 1991). Next at the Polar Front, it was found that local naturally elevated Fe concentrations made the difference in stimulating three distinct blooms of phytoplankton accompanied by uptake of CO2 (de Baar et al. 1995; Smetacek et al. 1997).

These and other findings led to a suite of large-scale in situ Fe fertilization experiments where an area of 50 km2 or even 200 km2 ocean surface water was fertilized with dissolved Fe. The major finding was that indeed, Fe addition did always stimulate phytoplankton blooms, but light limitation due to deep mixing in storms did also play a key role (de Baar et al. 2005). Therefore we now know that the phytoplankton in the HNLC regions is limited by a combination of light and iron limitation. The surface water values of Fe in the Antarctic (Fig. 4) are very low, notably ~0.1 nM near the Antarctic continent. There may be several reasons for this. Here the very deep (~1500 m) shelf under the continental ice sheet extending far over the sea is exceptional as it is not a source of Fe to overlying waters. The ice sheet acts as a cap largely preventing any biochemistry in the water column, and hence the underlying sediment is not a source of reduced Fe. In addition, the low dissolved Fe around Antarctica is due to very low dust input.

Panel B. Processes within ocean waters affecting the distribution of dissolved trace metals

Colloids . Throughout natural waters, both freshwater and seawater, there exists a continuum of physical/chemical state from truly dissolved to a wide range of colloidal sizes to truly particles. In the open ocean, colloids can be defined by ultrafiltration methods to be in the ~0.02–0.2 μm (micron) size classes (Nishioka et al. 2001, 2005), the pool <0.02 μm then is defined as the soluble fraction, and all sizes above 0.2 μm are defined as particles. Colloids can play a significant role, for example, formation of Fe colloids in hydrothermal plumes. However due to space restrictions, here we must largely ignore the colloids and operationally define the dissolved trace metals by the size cutoff of 0.2 μm pore size filtration devices.

Chemical speciation of dissolved trace metals results from the interaction of the free metal ion, for example, Fe3+ interacts with those ions of seawater to which it has great chemical affinity. For example, taking into account only the inorganic ions in seawater, the total dissolved Fe comprises 92% Fe(OH)3(dissolved), 4.5% Fe(OH)4 , 3.7% Fe(OH)2 +, and merely 3 × 10−9% free Fe3+ ion (Millero et al. 1995). Thus the concentration of free Fe3+ ion is a factor 3 × 10−11 less abundant than the total dissolved Fe(III). For the reduced Fe(II) state that can occur due to photoreduction (see below), there are 76% free Fe2+ ion, 23% FeCO3(dissolved), and minor concentrations of FeHCO3 +, Fe(CO3)2 2−, and Fe(OH)+ ions. Similar inorganic speciations have been calculated for all trace metal elements in seawater (Turner et al. 1981; Byrne et al. 1988). These calculations also include other divalent metals like Co(II), Ni(II), Cu(II), and Zn(II), trivalent metals like Al(III), and the REE(III), as well as for those metals existing as oxyanions such as V existing predominantly as vanadate HVO4 2− ion, Cr existing mostly as chromate CrO4 2− ion, and Mo existing mostly as molybdate MoO4 2− ion.

Organic complexation of the free ion by dissolved organic ligands (largely unknown functional groups) has been found to be very significant for several trace metals, most notably Fe, Zn, Cu, and Co, that also happen to be bio-essential trace elements, as well as Cd that shows a strong relationship with the biological cycle. Using voltametric techniques it has been found that ~99% of dissolved Fe is organically bound (Gledhill and Van den Berg 1994; Gledhill and Buck 2012). The above inorganic speciation thus represents merely ~1% of the total dissolved Fe. Similarly 98% of dissolved Zn is bound by organic complexes. Notably for Fe this organic complexation is favorable for maintaining dissolved Fe in solution; without the stabilization by organic complexation, most of the dissolved Fe would tend to be removed by oxidative precipitation.

Photoreduction within the upper euphotic zone of the oceans is known to cause a (diurnal) shift from the Fe(III) to the Fe(II) oxidation state, with implications for overall distribution of Fe in different pools, and hence the availability of Fe for uptake by biota (Rijkenberg et al. 2004, 2005, 2006). Similarly photoreduction of particulate MnO2 may lead to Mn dissolution such that the dissolved Mn2+ in surface seawater increases (Sunda et al. 1983; Sunda and Huntsman 1988, 1990, 1994) in favor of uptake by biota.

Uptake by phytoplankton (and also by bacteria and viruses) of bio-essential elements Mn, Fe, Co, Ni, Cu, and Zn is a major pathway in the ocean cycling of these trace metals, as per Eq. 2. Net result is the removal of dissolved metals from the euphotic zone (the upper water down to 1% penetration of incoming sunlight, typically 40–80 m depth in open oceans) in surface waters (all waters shallower than 100 m depth). The true internal metal content of phytoplankton has been difficult to distinguish from metals adsorbed at the outside of the cell; see below adsorptive scavenging. However novel irradiance techniques using a synchrotron have been a breakthrough in being able to clearly discern internal “true biological” metal atoms from those adsorbed external on the outside of the cell wall (Twining 2003; Twining and Baines 2013). Upon senescence by nutrient limitation and/or grazing or viral lysis, some “dead” biogenic debris settles down into intermediate waters (100–1000 m depth) where most is remineralized and the metals return to the dissolved state. Particles settling below 1000 m further mineralize and further return the metals to the dissolved state.

Adsorptive scavenging is the removal of dissolved trace metals from seawater by adsorption onto the outer surfaces of settling biogenic particles. This is very effective for adsorption of positive ions; the negatively charged oxyanions are not affected. Some redox metals, notably Mn and Fe, will next oxidize further at this surface, and such freshly formed Fe-Mn oxyhydroxides are known to be very effective adsorbing agents for a suite of other trace metals. Otherwise organic functional groups of the particles also are deemed to be effective scavengers (Balistrieri et al. 1981). Upon settling into intermediate (100–1000 m depth) and deep (>1000 m depth) waters, the biogenic debris is respired, and the trace metals are remineralized back again as dissolved metals into the seawater. Some trace metals immediately adsorb again onto other deep particles and are once again removed from solution. From concentrations of dissolved trace metals that remain in surface waters, it is impossible to derive how much has been removed by true internal biological uptake and how much removed by external adsorptive scavenging. However in the deep waters, one may discern deep water scavenging simply when the concentration of a bio-essential trace metal is less than expected from the concentration of a major nutrient phosphate or silicate. From such approach we know that there is significant deep water scavenging of Fe, Mn, Cu, and Ni, but apparently not of Zn and Cd.

Manganese (Mn) is essential in several metalloproteins within marine phytoplankton (Table 1). In seawater it exists in the reduced [Mn2+] state that in the well-oxygenated open ocean is prone to slow oxidation to the insoluble MnO2. However in reducing sediments of the continental shelf, the MnO2 dissolves again to high levels of [Mn2+] in pore waters, and the [Mn2+] diffuses into overlying waters, where it can be traced over long distances (Wu et al. 2014). In an oxygen minimum zone, this higher [Mn2+] can be mixed over even longer lateral distances, e.g., in the East Equatorial Pacific (Martin and Knauer 1985) and the Northwest Indian Ocean (Saager et al. 1989).

In the open ocean, the surface water concentrations of dissolved Mn can range up to ~2–4 nM due to partial dissolution of aeolian dust supply from land. Photoreduction of Mn in the euphotic zone is known to enhance such dissolution (Sunda et al. 1983; Sunda and Huntsman 1988, 1990, 1994). The ~2–4 nM concentrations are deemed to be in ample supply for the needs of phytoplankton. In surface waters of the remote Southern Ocean with low dust input, the dissolved Mn is considerably lower in the 0.1–0.9 nM range (Fig. 4; Middag et al. 2011b, 2012, 2013). The lowest values in the 0.04–0.12 nM range are near the Antarctic continent, where the very deep (~1500 m) shelf under the extended continental ice sheet is exceptional as it is not a source of Mn (Middag et al. 2011b), as opposed to the extensive shelves around the Antarctic Peninsula acting as a source of Mn (Fig. 3) due to reductive dissolution in the anoxic pore waters (Middag et al. 2012, 2013).

In Antarctic surface waters, the net removal of dissolved Mn covaries with net removal of part of the dissolved phosphate, with an apparent Mn:P uptake ratio by plankton in the 0.36–0.39 × 10−3 mol.mol−1 range (Middag et al. 2011b, 2013). In the deep waters of Atlantic and Antarctic Oceans, the dissolved Mn is remarkably uniform at very low 0.1–0.15 nM concentrations (Van Hulten et al. 2017). However, Mn plumes with much higher concentrations up to 3 nM or even 10 nM (Arctic Gakkel Ridge) exist over mid-ocean ridges due to hydrothermal input, the plumes extending several hundred kilometers horizontally and almost 1000 m above the ridge crest (Middag et al. 2011a, b; Resing et al. 2015, Wu et al. 2014). Sampling with submersibles of the emitting high temperature strongly reducing vent fluid gives very high Mn concentrations up to ~1 millimolar concentration, i.e., up to ten million times the low background dissolved concentration of ~0.1 nM (Von Damm et al. 1985, 1998 versus Middag et al. 2011b, 2013; Van Hulten et al. 2017). Just like for Fe (see above), most of the hydrothermal Mn deposits rapidly in metalliferous sediments at the ridge crest, yet the million times lower 1–10 nM concentration remaining in the plume can be seen over more than 1000 km away from the ridge crest, versus the low 0.1 nM background concentration in deep waters (Fig. 4).

Zinc (Zn) is the second most important trace metal element in biology. Its most important function is in the enzyme carbonic anhydrase that enhances the reaction rate some ten million times for the conversion of bicarbonate HCO3 ion to CO2. For phytoplankton this is most helpful as it allows to take up not only CO2 but also HCO3 , which is far more abundant, and convert the latter to CO2, which is necessary as only CO2 is utilized in photosynthesis. Thus at conditions of relatively low ambient [CO2]aqueous, the algae can also assimilate HCO3 .

The divalent Zn(II) state in seawater is ~98% bound by organic complexation with dissolved organic molecules; the remaining is ~1.5% free [Zn2+] ion and ~0.5% inorganic species [ZnCO3(aqueous)] and [ZnSO4(aqueous)]. Presumably the 98% organic complexation helps to maintain the Zn in solution because the free [Zn2+] ion is prone to loss due to adsorptive scavenging on marine particles.

The ocean distribution of dissolved Zn very closely follows that of major nutrient silicate (Fig. 5). This is consistent with its biochemical functions (Table 1), and perhaps the close relation with silicate hints at a key biological relationship with diatoms. Surface water concentrations are low in the order of 0.1 nM but higher at 2–3 nM in the Antarctic Ocean, such that this is not only a HNLC region for major nutrients but also for trace nutrient Zn. Deep water concentrations increase from 2 nM in the deep North Atlantic to 6–8 nM in the deep Antarctic to 10 nM in the deep North Pacific.

Copper (Cu) has several functions in biology (Table 1). In seawater the divalent Cu(II) state is virtually 100% bound by strong organic complexation. There is ample total dissolved Cu at ~0.5–1.5 nM in surface waters of the North Atlantic, and therefore Cu is deemed not to be limiting for phytoplankton growth, unless the very strong organic complexation would hinder plankton uptake. The concentration increases steadily with depth to 2.0–2.5 nM at around 4000 m depth. This steady increase resembles that of major nutrient silicate. Nevertheless when comparing the deep North Atlantic with the deep North Pacific, the steady increase of dissolved silicate with increasing age of deep water is hardly followed by an increase of dissolved Cu in older deep waters of the Pacific. Therefore it is generally assumed that there is net scavenging removal of Cu from deep waters that prevents Cu from accumulating a high dissolved concentration in old deep waters.

Nickel (Ni) has perhaps the least biochemical functions of the bio-essential trace elements (Table 1). In seawater Ni dominantly exists as the free hydrated Ni(II) state. The vertical distribution of Ni is perhaps best described by a correlation with both phosphate and silicate. There is ample Ni at ~2 nM in surface waters; therefore Ni is deemed not to be a limiting element for phytoplankton growth.

Cobalt (Co) has a key functionality as the metallocenter of vitamin B12 and can also substitute for Zn in carbonic anhydrase. Cobalt exists in the Co(II) state in seawater and is for ~90% bound as Co-organic complexes. Cobalt also has similar redox chemistry to that of iron and manganese. The oxidized Co(III) state is highly insoluble due to precipitation of Co(III) oxyhydroxides. Conversely, at low ambient O2, this may be reduced again, and Co(II) is remobilized into the dissolved state. Indeed concentrations of Co, Fe, and Mn are considerably higher in O2 minimum zones, such as off West Africa (Noble et al. 2012).

Vanadium (V) exists in seawater primarily as the vanadate oxyanion [HVO4 2−] (Butler 1990) at a concentration between 30 and 36 nmol.kg−1 (Collier 1984). The role of V in proteins found in blood cells of ascidians (sea squirts) is very special (Michibata and Sakurai 1990) and most intriguing but so unique that it does not affect the concentration of V in ocean waters.

Molybdenum (Mo) is together with Fe required for the process of N2 fixation by the so-called diazotrophs (N2 fixers) among which Trichodesmium is the “model” organism (La Roche and Breitbarth 2005). Molybdenum exists in seawater as Mo(IV) in the molybdate oxyanion state [MoO4 2−] and as such not prone to adsorptive scavenging loss. The high concentration in seawater at ~104 nM renders Mo the most abundant trace metal element in the oceans, where it is deemed a conservative element varying only slightly as function of salinity (Collier 1985). Given this high abundance, Mo is by no means limiting for growth and functioning of the diazotrophs; however, the other trace element Fe as required for N2 fixation is often very low and limiting in surface waters. The process of N2 fixation also requires high seawater temperatures above ~20 °C. In the warm waters of the Sargasso Sea, directly under the influence of major deposition of Fe-rich aeolian dust derived from the Sahara, there exists an important niche for Trichodesmium to form major blooms. Recently the requirement of temperatures above ~20 °C has been challenged (Rivero-Calle et al. 2016).

Cadmium (Cd) was long considered to have no biological function, at most toxic, as it may as a look-alike of Zn occupy, and hence disable, places in the cell where Zn ought to be and function. However the discovery of the very close relationship of dissolved Cd with major nutrient phosphate in the oceans (Boyle et al. 1976; Bruland 1980) led to more investigations of interactions between Cd and phytoplankton. Now we know that in fact Cd can substitute for Zn in phytoplankton, in conditions where there is not enough Zn available. Moreover there exist intrinsic Cd-cofactored carbonic anhydrases, for certain diatoms thus far studied, i.e., sometime during biological evolution, this true Cd-based enzyme has developed (Lane and Morel 2000; Lane et al. 2005; Xu et al. 2008). In seawater the Cd(II) exists for some 64% as an organic complex, the remaining Cd(II) exists for some 30% as chlorine inorganic species, and merely 3% free [Cd2+] ion.

In general surface water concentrations of Cd are very low indicating uptake and/or scavenging removal by plankton. However surface waters of the Southern Ocean have high dissolved Cd, thus here also mimicking the HNLC condition of phosphate. The North Atlantic Deep Water (NADW) originates from Atlantic surface waters and has relatively low concentrations of both Cd and phosphate, as opposed to the high concentrations in the Antarctic Intermediate Water (AAIW) and Antarctic Bottom Water (AABW) coming from the south and flowing northward.

An Update on the Extended Redfield Stoichiometry

The stoichiometry of the nine bio-essential major and trace elements most likely varies considerably between phytoplankton species, as well as due to some degree of adjustment that the algae can make vis-à-vis availability of each of these nine essential elements in the ambient seawater. Conversely depending on the essential elements being more or less available, the composition of the plankton will vary, albeit within certain limits below which the cell cannot function anymore.

During more than 25 years, it has been, and still very much is, a challenge to resolve the stoichiometric “constants” of the six bio-essential trace metals (e.g., Bruland et al. 1991; De Baar et al. 2008). In an excellent recent review by Twining and Baines (2013), the various approaches have been synthesized. From this the order of magnitude values are listed in Table 3, by no means as the final truth but merely as an indication of the current state-of-the-art.
Table 2

Concentrations in seawater of the nine chemical elements that are essential for all life in bold print as well as four chemical elements that play a role in specific taxonomic groups. Each element symbol is preceded by its atomic number #. Major dissolved inorganic species, i.e., not taking into account here the strong organic complexation for several trace elements as discussed in main text for such element. Given concentrations are typical values; real concentrations vary around these as function of depth and geographic location. Not listed are Mg and Ca that also play a role in specific taxonomic groups but are listed among the major ions of sea salt (see entry Ocean Salinity, Major Elements, and Thermohaline Circulation), as is S in abundant dissolved sulfate, and hence Mg, Ca, and S never are bio-limiting

#

 

Major inorganic species

Concentration

North Atlantic Ocean

Antarctic Ocean

North Pacific Ocean

Unit

Surface

Deep

Surface

Deep

Surface

Deep

6

C

HCO3

[μmol.kg−1]

2050

2200

2210–2220

2240–2260

2000

~2350

7

N

NO3

[μmol.kg−1]

<1

25

25–30

30–38

<1

45

14

Si

H4SiO4

[μmol.kg−1]

<1

40

35–63

83–129

<1

170

15

P

HPO4

[μmol.kg−1]

<0.1

1.2–1.5

1.6–1.9

2.26–2.35

<0.1

2.5–3.3

23

V

HVO4 2−

[nmol.kg−1]

Quite uniform at ~30–37 [μmol.kg−1 in all ocean waters, mostly conservative

25

Mn

Mn2+

[nmol.kg−1]

0.2–3.0

0.1–0.15

0.04–0.5

~0.1

~1–2

~0.2

26

Fe

Fe(OH)2+

[nmol.kg−1]

0.2–2.0

0.6

0.02–0.1

0.4

0.05

0.6

27

Co

Co2+

[pmol.kg−1]

30

60

20–85

~60

10–50

50–250

28

Ni

Ni2+

[nmol.kg−1]

2

4

5.5–6.5

5.5–7.0

2.1

11

29

Cu

CuCO3 0

[nmol.kg−1]

1 ± 0.5

1.3–2.2

1–2

2–4

<0.5

5

30

Zn

Zn2+

[nmol.kg−1]

0.1

2

2.5–3.1

6.2–7.8

0.07

9

42

Mo

MoO4 2−

[nmol.kg−1]

Quite uniform at ~104 μmol.kg−1 in oceans, conservative varying slightly with salinity

48

Cd

CdCl2 0

[pmol.kg−1]

10

300

400–600

770–860

1.4

1100

DIC Atlantic: see Fig. 4; DIC Antarctic: van Heuven et al. 2011; DIC Pacific: see Fig. 4.

P Atlantic: see Fig. 4; P Antarctic: Baars et al. (2014); P Pacific: see Fig. 4.

N in Atlantic: see Fig. 4; N in Antarctic: de Baar et al. 1997; N in Pacific: see Fig. 4.

Si Atlantic: see Fig. 4; Si Antarctic: Zhao et al. 2014; Si Pacific: see Fig. 4.

V worldwide: Collier (1984).

Mn Atlantic: Van Hulten et al. 2017; Mn Antarctic: Middag et al. 2011b, 2012, 2013; Mn Pacific: Resing et al. 2015.

Fe Atlantic: Rijkenberg et al. 2014; Fe Antarctic: Klunder et al. 2011, 2014 and Gerringa et al. 2015; Fe Pacific: Martin et al. 1989.

Co Atlantic: Middag et al. 2015b, Noble et al. 2012. Co Antarctic: Saito et al. 2010; Co Pacific: Hawco et al. 2016, values >50 pM in O2 minimum zone.

Ni Atlantic: Middag et al. 2015b; Ni Antarctic: Cameron and Vance 2014; Ni Pacific: Bruland 1980.

Cu Atlantic: Middag et al. 2015b; Cu Antarctic; Bown et al. 2016; Cu Pacific: Bruland 1980.

Zn Atlantic: Middag et al. 2015b; Zn Antarctic: Zhao et al. (2014); Zn Pacific: Martin et al. 1989.

Mo worldwide: Collier (1985); see also Singh et al. (2011).

Cd Atlantic: Middag et al. (2015b); Cd Antarctic: Baars et al. (2014); Cd Pacific: Bruland 1980.

Data tables in Bruland (1980), Martin et al. (1989), Baars et al. (2014), and Zhao et al. (2014) and in online supplement of Gerringa et al. (2015). The other datasets available at www.geotraces.org and in the datasets cited in captions of Figs. 2, 3, 4, 5, 6, 7, and 8.

Table 3

An illustration of the current (2017) state-of-the-art for the order of magnitude of stoichiometric “constants” for the extended Redfield Eq. 2. As per the latter Eq. 2, the data are normalized to value 1 for P (phosphorus). Values for major elements C, N, P, and O2 after Anderson and Sarmiento (1994). Instead for the trace elements Fe, Zn, Mn, Cu, Ni, and Co, typical ranges are given after Twining and Baines (2013; their Fig. 4)

Major elements

C

(117 ± 14)

N

(16 ± 1)

P

1

O2

(−170 ± 10)

Trace elements

Fe

(0.2–20) × 10−3

Zn

(0.1–0.6) × 10−3

Mn

(0.2–4.0) × 10−3

Cu

(0.2–2.0) × 10−3

Ni

(0.1–1.8) × 10−3

Co

(0.005–2.0) × 10−3

In general the mixing of ocean waters tends to some extent to “average” the wide variability in living plankton and settling biogenic debris. Nevertheless the distributions of each of the nine bio-essential elements in the oceans have more variability than most people realize, this also for the “classical” major elements. Therefore upon further investigations, it remains to be seen whether or not the ranges as listed in Table 3 (e.g., 117 ± 14 for carbon or (10 ± 9) × 10−3 for iron) will in due course become either narrower or wider.

Recently completed and ongoing world ocean research programs such as WOCE, CLIVAR, and GEOTRACES (www.geotraces.org) are providing newer, larger, and more accurate datasets of the dissolved major and trace elements, which undoubtedly will lead to future adjustments and refinements in the numbers presented in Table 3. For example, the apparent biological uptake ratio Mn/P derived from water column profiles of GEOTRACES in 2008 in the Southern Ocean is in the 0.3–0.4 × 10−3 range (Middag et al. 2013), that is, in the lower part of the range in Table 3.

Other Trace Elements

Recently Bruland et al. (2014) produced an excellent overview of the trace metals in seawater that comprises all trace elements, those with key biological functions as above discussed and virtually all of those, more than 40 elements, that are not bio-limiting, their abundance not controlled by biology, or without a biological function at all. They also provide an extensive table comprising almost all trace elements and their abundance in seawater. Below only 3 of the 55 nonbiological trace elements are discussed, but Table 4 comprises all 55 largely abiotic trace elements, complementary to Tables 1 and 2 for “biological” elements.
Table 4

Concentrations in seawater of the 55 trace elements that are not bio-limiting, and/or their abundance not controlled by biology, or without a biological function at all. For elements with key biological functions, see Tables 1 and 2. Not listed are noble gases and elements of which no stable isotopes exist in seawater, although radioisotopes 232Th and 238U with very long half-life are listed; the long half-life is such that these are treated as virtually stable isotopes in the oceans. This Table 4 is based on the informative Table in Bruland et al. (2014), plus a few extra chemical elements not listed by Bruland et al. (2014)

Atomic number

Symbol

Name

Major species

Concentration range

Concentration typical

3

Li

Lithium

Li+

(Conservative)

25.9 μmol.kg1

4

Be

Beryllium

Be(OH)+, Be(OH)2 0

4–30 pmol.kg−1

~23 pmol.kg−1

5

B

Boron

H3BO3

(Conservative)

416 μmol.kg−1

13

Al

Aluminum

Al(OH)4 , Al(OH)3 0

0.3–40 nmol.kg−1

~2 nmol.kg−1

21

Sc

Scandium

Sc(OH)3 0

8–20 pmol.kg−1

~16 pmol.kg−1

22

Ti

Titanium

Ti(OH)4 0, TiO(OH)2 0

6–250 pmol.kg−1

~150 pmol.kg1

24

Cr

Chromium

CrO4

3–5 nmol.kg−1

~4 nmol.kg−1

31

Ga

Gallium

Ga(OH)4

5–60 pmol.kg−1

~20 pmol.kg−1

32

Ge

Germanium

H4GeO4

1–100 pmol.kg−1

~70 pmol.kg−1

32

Ge

MethylGe

CH3Ge(OH)3 0, (CH3)2, Ge(OH)0

(Two conservative species)

400 pmol.kg−1

33

As

Arsenic

HAsO4 2−

17–25 nmol.kg−1

23 nmol.kg−1

34

Se

Selenium

SeO4 2−

0.5–2.3 nmol.kg−1

1.8 nmol.kg−1

37

Rb

Rubidium

Rb+

(Conservative)

~1.4 μmol.kg−1

39

Y

Yttrium

Y(CO3)+, Y(oh)2 +

60–300 pmol.kg−1

200 pmol.kg−1

40

Zr

Zirconium

Zr(OH)5 , Zr(OH)4 0

9–300 pmol.kg−1

~160 pmol.kg−1

41

Nb

Niobium

Nb(OH)6 , Nb(OH)5 0

1–4 pmol.kg−1

~ 3 pmol.kg−1

43

(Tc)

(Technetium)

(Does not occur in nature, except extinct Oklo natural nuclear fission reactor in Gabon)

44

Ru

Ruthenium

Ru(OH)n 4-n

 

<50 fmol.kg−1?

45

Rh

Rhenium

Rh(OH)n 3-n, Rh(cl)n 3-n

0.4–1 pmol.kg−1

0.8 pmol.kg−1

46

Pd

Palladium

PdCl4

0.2–1 pmol.kg−1

~0.7 pmol.kg−1

47

Ag

Silver

AgCl2 , AgCl3 2−

1–35 pmol.kg−1

20 pmol.kg−1

49

In

Indium

In(OH)3 0

0.04–2 pmol.kg−1

~0.1 pmol.kg−1

50

Sn

Tin

SnO(OH)3 , Sn(OH)4 0

1–20 pmol.kg−1

4 pmol.kg−1?

51

Sb

Antimony

Sb(OH)6

(~Conservative)

1.6 nmol.kg−1

52

Te

Tellurium

Te(OH)6 0

0.5–1.5 pmol.kg−1

~0.6 pmol.kg−1

53

I

Iodine

IO3

350–460 nmol.kg−1

450 nmol.kg−1

55

Cs

Cesium

Cs+

(Conservative)

2.2 nmol.kg−1

56

Ba

Barium

Ba2+

30–150 nmol.kg−1

110 nmol.kg−1

57

La

Lanthanum

LaCO3 +

8–40 pmol.kg−1

~35 pmol.kg−1

58

Ce

Cerium

CeCO3 +

1.5–10 pmol.kg−1

5 pmol.kg−1

59

Pr

Praseodymium

PrCO3 +

1–8 pmol.kg−1

4 pmol.kg−1

60

Nd

Neodymium

NdCO3 +

4–50 pmol.kg−1

30 pmol.kg−1

61

(Pm)

(Promethium)

Does not occur in nature (unstable radioisotopes can form in nuclear reactor)

62

Sm

Samarium

SmCO3 +

1–10 pmol.kg−1

5 pmol.kg−1

63

Eu

Europium

EuCO3 +

0.2–2.5 pmol.kg−1

0.8 pmol.kg−1

64

Gd

Gadolinium

Gd(CO3)2−

1.5–12 pmol.kg−1

9 pmol.kg−1

65

Tb

Terbium

Tb(CO3)2−

0.5–2 pmol.kg−1

~1 pmol.kg−1

66

Dy

Dysprosium

Dy(CO3)2−

2–14 pmol.kg−1

9 pmol.kg−1

67

Ho

Holmium

Ho(CO3)2−

0.8–2.2 pmol.kg−1

1.6 pmol.kg−1

68

Er

Erbium

Er(CO3)2−

1.5–12 pmol.kg−1

10 pmol.kg−1

69

Tm

Thulium

Tm(CO3)2−

0.3–2 pmol.kg−1

1 pmol.kg−1

70

Yb

Ytterbium

Yb(CO3)2−

1–12 pmol.kg−1

6 pmol.kg−1

71

Lu

Lutetium

Lu(CO3)2−

0.2–2 pmol.kg−1

~1 pmol.kg−1

72

Hf

Hafnium

Hf(OH)4 0, Hf(OH)5

0.06–1 pmol.kg−1

~0.3 pmol.kg−1

73

Ta

Tantalum

Ta(OH)5 0

0.01–0.3 pmol.kg−1

~0.1 pmol.kg−1

74

W

Tungsten

WO4 2−

(Conservative)

55 pmol.kg−1

75

Re

Rhenium

ReO4

(Conservative)

40 pmol.kg−1

76

Os

Osmium

H3OsO6

3–8 fmol.kg−1

~ 6 fmol.kg−1

77

Ir

Iridium

Ir(OH)3 0 (?)

0.5–1 fmol.kg−1

 

78

Pt

Platinum

PtCl4 2−?

0.2–1.5 pmol.kg−1

~0.8 pmol.kg−1?

79

Au

Gold

AuCl2 ?, Au(OH)3 0?

10–100 fmol.kg−1

~50 fmol.kg−1

80

Hg

Mercury

HgCl4 2−

0.2–10 pmol.kg−1

~1 pmol.kg−1?

81

Tl

Thallium

Tl+, TlCl0

60–75 pmol.kg−1

70 pmol.kg−1

82

Pb

Lead

PbCO3 0

4–150 pmol.kg−1

~10 pmol.kg−1

83

Bi

Bismuth

Bi(OH)2+, BiO+

0.04–0.5 pmol.kg−1

~0.1 pmol.kg−1

90

232Th

Thorium

Th(OH)3(CO3)?

0.3–0.6 pmol.kg−1

~0.3 pmol.kg−1

92

238U

Uranium

UO2(CO3)3 4−

(Conservative)

~13.5 nmol.kg−1

Aluminum (Al) has a high abundance in rocks and soils of the planet but has very low concentrations in the oceans. Dissolved trivalent Al is highly prone to adsorption on particles and perhaps the ultimate scavenging-type trace element in the oceans. The major external source for Al is deposition of aeolian dust coming from the continents. Consequently, an elevated concentration of dissolved Al in surface seawater is strong evidence for dust input, notably Sahara dust that is blown by prevailing winds all across the North Atlantic in the 10–20 °N region (Fig. 6). However due to the intensive scavenging loss, Al rapidly disappears from seawater. This is nicely shown in Fig. 7 where at first glance the distribution of Al looks like that of major nutrient silicate, yet in fact it is the opposite. Where silicate is low, Al is high; where Si is high, Al is low. This amazing mirror image of Al versus Si is partly due to North Atlantic surface water being depleted in nutrient Si but receiving massive supply of Al in aeolian dust from the Sahara.
Fig. 7

Section West Atlantic, Geotraces GA02. Upper graph: silicate [μmol/kg], after Middag et al. (2015a). Middle graph: Zn [nM], after Middag et al. (2015b). Lower graph: Al [nM] after Middag et al. (2015a). Vertical axis is pressure [decibar] that is virtually identical to depth [meter] (All data as shown here by Middag and coworkers is in Geotraces IDP 2014 (www.geotraces.org). Graphics www.eGeotraces.org using Ocean Data View (R. Schlitzer))

In contrast, in the Antarctic Ocean, the surface silicate is high (HNCL region), but with the Antarctic being so remote from land, there is no significant dust input of Al. As a result the Antarctic waters are high in Si and very low in Al (Middag et al. 2011c, 2012, 2013). When these Antarctic waters flow northward as AAIW and AABW into the Atlantic basin, they bring along the high Si but low Al signal. This in combination with intensive scavenging may explain the opposite distributions. Indeed by ocean biogeochemical simulation modeling of these processes, it is possible to reproduce the distributions of both Si and Al along this section (Van Hulten et al. 2014).

Iridium (Ir) has the record of lowest concentration of trace elements in the oceans. Iridium is extremely rare in the Earth’s crust. This in fact is the reason why an anomalous Ir-enriched layer in sedimentary rocks was identified as clear evidence of asteroid impact, asteroids having a higher Ir concentration than crustal rocks of planet Earth (Smit and Hertogen 1980; Alvarez et al. 1980). This asteroid impact is deemed to have caused the demise of the dinosaurs. Iridium is trivalent in seawater and akin to above trivalent Al likely to be also rapidly scavenged. Thus any input from land that in the first place already comprises very little Ir would next also be removed quickly from the ocean water column. Indeed dissolved Ir is extremely low in the 0.5–1.0 femtomolar range (1 femtomol.L−1 = fM = 10−15 mol.L−1) (Anbar et al. 1996).

Lead has a very low natural background dissolved concentration (Pb<10 pmol), and this is consistent with the fact that it is very susceptible to adsorptive scavenging. However due to the use from ~1930 onward of gasoline with tetraethyl-Pb additive (as an antiknock agent), massive amounts of Pb have been emitted in the atmosphere over North America and Europe. This large-scale pollution was noticed by Claire Patterson through his efforts to date the age of our planet using Pb isotopic abundances in rocks, which occasionally led to spurious results. Patterson realized this was due to inadvertent contamination in his laboratory and was one of the first geochemists to construct a clean room laboratory and to develop and apply rigorously ultraclean methods. This was successful, and eventually in 1956, the date of the Earth was estimated at 4.55 billion years, a value repeatedly confirmed.

Meanwhile Patterson initiated pioneering work to detect traces of Pb in the environment. Using the best ultraclean sampling methods of the time, samples were collected and analyzed for dissolved Pb in the Pacific Ocean and in year 1979 in the Sargasso Sea near Bermuda in the North Atlantic Ocean (Flegal and Patterson 1983; Schaule and Patterson 1981, 1983). Due to prevailing wind directions, the North Atlantic is downwind of North America and major recipient of the anthropogenic Pb signal.

Moreover, Patterson almost singlehandedly crusaded to abandon Pb in gasoline and eventually succeeded, such that the emissions of Pb into the atmosphere have greatly diminished from the late 1970s onward (Bryson 2003). The Bermuda Atlantic Time-Series Station (BATS) showed that concentrations of dissolved Pb steadily decreased from maximum ~160–170 pM in the upper 500 m in 1979 (Schaule and Patterson 1983) to ~20 pM in the same 500 m depth range in year 2011(Boyle et al. 2014).

Using the most advanced state-of-the-art clean PRISTINE sampling equipment (Rijkenberg et al. 2015) along the complete West Atlantic section (including the BATS site), the dissolved Pb was measured (Fig. 8).
Fig. 8

Section of dissolved Pb in the West Atlantic Ocean, Geotraces. Vertical axis is pressure [decibar] that is virtually identical to depth [meter]. The still contaminated North Atlantic waters and overall surface waters are in contrast with the subsurface waters of Antarctic origin (AAIW at ~1000 m and AABW over the seafloor) that are virtually free of Pb. The vertical profile of this dataset at the BATS near Bermuda is reported by Middag et al. (2015) (See also Malakoff 2014). All data as shown here by Middag and coworkers is in Geotraces IDP 2014 (www.geotraces.org). Graphics www.eGeotraces.org using Ocean Data View (R. Schlitzer)

The upper waters of the North Atlantic nowadays have lower dissolved Pb than in the 1979–2008 era, but the subsurface waters in the ~500 to ~2500 m depth range still carry part of the pollutant signal of past high Pb emissions. Given the scavenging-type nature of dissolved Pb, luckily this pollutant signal will eventually disappear from the ocean water column. In the southernmost part of the section, it is noticed that the subsurface waters of Antarctic origin (AAIW at ~1000 m and AABW over the seafloor) are virtually free of Pb, apparently truly “pristine,” due to the fact that the dust supply to the Antarctic Ocean is very low.

Combination of this long section (Fig. 8) with two meridional sections in the North and South Atlantic, by the USA and the UK, respectively, has been shown and discussed by Malakoff (2014).

Summary

The natural chemical elements exist in seawater in concentrations ranging over 15 orders of magnitude from Cl at ~0.56 mol.kg−1 seawater down to Ir at ~0.5 × 10−15 mol.kg−1 seawater. The most abundant ions exist in constant proportions. Together these major ions constitute the dissolved salt content or salinity of seawater.

Life in the oceans requires not only C, N, and P but also six trace nutrient elements Mn, Fe, Co, Ni, Cu, and Zn. Photosynthesis by marine algae utilizes these overall nine nutrient elements from seawater in fairly uniform proportions or stoichiometry. The reverse remineralization or respiration by marine bacteria and animals recycles these nutrient elements back into the seawater. Moreover important taxonomic groups also use other chemical elements, most notably Ca for producing CaCO3 shells and coral reefs and Si for producing opaline SiO2 external frustules of diatom algae. In the temperate and tropical oceans, the major nutrients N, P, and Si are depleted from surface waters. In contrast in the Antarctic Ocean and some other high-nutrient low-chlorophyll (HNLC) regions, there are plenty of major nutrients N, P, and Si, but the ecosystem is limited due to a lack of Fe in combination with light limitation. Due to deep remineralization, there are higher concentrations of nutrients N, P, Si, Co, Ni, Cu, and Zn in deep waters. The trace nutrients Mn and Fe have multiple oxidation states. Both Mn and Fe are involved in the biological cycle but also are affected by aeolian dust supply, oxidative scavenging removal from seawater, and opposed reductive dissolution within anoxic marine sediments and supply from hydrothermal vents.

The remaining 55 chemical elements and their low concentrations in seawater are listed. Yet only three are discussed: Al is the key tracer for aeolian dust input into the oceans, Ir has the record of lowest concentration in seawater, and Pb has elevated abundance due to past use of leaded gasoline.

Cross-References

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Copyright information

© Springer International Publishing AG 2017

Authors and Affiliations

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