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5 Light, Nutrients and Oxygen in the Sea

5 Light, Nutrients and Oxygen in the Sea

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12



1 Preliminaries



eventually only blue or violet light is left. In the clearest ocean, light intensity is

reduced to 1 % of the surface value at a depth of *50–100 m, but this intensity can

be reached at depths as shallow as 5–10 m in turbid coastal waters, where high

levels of suspended material reduce light penetration. The 1 % level, called

euphotic depth, is often used to define the base of the euphotic zone, the depth

below which light intensity is too low for photosynthesis to take place.

Mathematically, vertical attenuation of light intensity (I) with depth (z) can be

described by the exponential equation:

Izị ẳ Io expðÀkzÞ;



ð1:2Þ



where Io is the light intensity at the sea surface, and k is the attenuation coefficient,

which depends strongly on turbidity (i.e. the concentration of particulate matter in

the water column) (Fig. 1.6). Photosynthesis depends on the availability of light

energy. Hence, the rate at which primary production of organic carbon can occur

also decreases rapidly with depth from a maximum at the ocean surface. However,

the rate of metabolic energy use by phytoplankton (respiration) varies little with

depth, leading to the concept of a compensation depth—the depth at which phytoplankton production equals respiration (Gran and Trygve 1935). Below this

depth, net phytoplankton production is not possible as respiration dominates over

production. Essentially, the compensation depth is similar to the euphotic depth.



Fig. 1.6 Relation between daily averaged light intensity and production/respiration of organic

carbon. The compensation depth is the depth at which production and respiration rate are equal.

Adapted from Segar (2007)



1.5 Light, Nutrients and Oxygen in the Sea



1.5.3



13



Oxygen



The availability of dissolved oxygen (O2) is vital for almost all marine life, with the

exceptions of either bacteria that can use sulphur or methane or their associated vent

communities. Because there is a continuous air-sea flux of gases between the

atmosphere and ocean, there is typically plenty of dissolved oxygen in surface

waters to support marine life. Below the surface in the ocean’s interior, however,

the oxygen level within a particular water mass depends on when the water was last

at the surface, how much oxygen has been used up during remineralization of

detritus (dead organic matter) as the water mass has moved along its path, and how

much it has mixed with other water masses during its journey. Hence, the dissolved

oxygen level at a given location depends on both ventilation age (time elapsed since

the last contact of a water mass with the atmosphere) and oxygen utilization.

Temperature is also important, as cold water can hold more dissolved oxygen than

hot water.

As has long been understood, the combination of significant utilization and weak

ventilation leads to a mid-depth oxygen minimum (Sverdrup 1938; Wyrtki 1962;

Fig. 1.7). Shallower waters tend to have higher dissolved oxygen concentrations

because they are better ventilated, despite higher rates of utilization, while deeper

waters (below about 2,000 m depth) tend to contain more oxygen because of lower

utilization rates and higher initial oxygen levels when they leave the surface, as they

are derived from cold, high-latitude source waters in e.g., the Greenland and

Norwegian Seas or around Antarctica, despite also having longer ventilation ages.

In most oceanic regions, oxygen-minimum layers are typically located at depths

of between 400 and 1200 m near the base of the permanent thermocline (Keeling

et al. 2010). In some regions however, such as the Northern Indian Ocean and most

of the Pacific Ocean, the oxygen minimum is associated with critically low,

potentially lethal oxygen levels. While these layers are typically located well below



Fig. 1.7 Typical vertical profiles of a seawater density, b dissolved oxygen, and c nitrate nitrogen

in the central Atlantic Ocean. Adapted from Segar (2007)



14



1 Preliminaries



the euphotic zone in most regions of the oceans, they can come close to the sea

surface (depths of 150–300 m) in the vicinity of the major coastal upwelling

regions of the tropical eastern Pacific and Atlantic Oceans and may even rise to

affect the euphotic zone (see Chap. 4).

The sensitivity of organisms, particularly macro-organisms, to changes in oxygen levels is variable. Most organisms are not very sensitive to oxygen levels as

long as the concentrations remain above a certain threshold, but once the oxygen

concentration falls below this threshold, the organism suffers from a variety of

stresses, leading ultimately to death if the concentration stays too low for too long.

Such low oxygen conditions are termed hypoxic or anoxic depending on whether

the oxygen concentration is merely low or totally absent, and thresholds for hypoxia

vary greatly between marine taxa, with fish and crustaceans tending to be the most

sensitive (Fig. 1.8).

A typical threshold for hypoxia is *60 lM/kg dissolved oxygen (Gray et al.

2002), which is equivalent to *1.4 mL/L or 2 mg/L. Hypoxic zones are often

called dead zones as the low oxygen concentrations can be lethal for many marine

animals, particularly benthic or burrowing organisms that cannot move rapidly, and

the area of coastal waters that is affected by hypoxia is increasing because of the

runoff of nutrients from farming and other anthropogenic inputs (Breitburg et al.

2009; Diaz and Rosenberg 2008). The added nutrients fuel increased phytoplankton

production, and when the phytoplankton die, their cells sink to the bottom and



Fig. 1.8 Median lethal oxygen concentration for four different taxa. Boxes run from the lower

(25 %) to the upper (75 %) quartile and also include the median. Redrawn after Vaquer-Sunyer

and Duarte (2008). Crustacea form a very large group of arthropods which includes familiar

animals such as crabs, lobsters, crayfish, shrimp, krill and barnacles. The species range in size from

Stygotantulus stocki at 0.1 mm to the Japanese spider crab with a leg span of up to 3.8 m. Bivalva

is a class of marine molluscs including clams, oysters, cockles, mussels and scallops. Gastropoda

(gastropods) include snails and slugs of a large range of sizes



1.5 Light, Nutrients and Oxygen in the Sea



15



decompose, using up dissolved oxygen. This is particularly important in stratified

water bodies, where there is a sudden change in density, as the resulting pycnocline

prevents oxygen from mixing downwards from the upper, oxygen-rich layer into

the bottom layer where decomposition is occurring. Hypoxic conditions, which may

be permanent or temporary, depending on local mixing conditions, are typical of

coastal upwelling regions (e.g, Monteiro et al. 2006).



1.5.4



Nutrients



A suite of chemicals, typically identified as nutrients, is required for phytoplankton

production in the ocean. Broadly important nutrients include nitrogen (N), phosphorus (P), silicon (Si) and iron (Fe). All phytoplankton taxa have relatively uniform requirements for N and P, but only diatoms and radiolarian protozoa require

silica. Plankton build their biomass with C:N:P ratios (C is carbon) of *106:16:1,

known as the Redfield ratio (Redfield 1958). Due to their high abundance in seawater, carbon and calcium, which is needed by the many organisms (including

corals and crustaceans) that create calcium carbonate shells, are typically not listed

among the nutrient elements.

Nutrient concentrations are generally low in the euphotic zone, because of their

uptake during photosynthesis, high rates of biological utilization and gravitational

settling of detritus. The remineralization of detritus as it sinks to greater depths

brings nutrients back into solution (and uses up dissolved oxygen in the process).

Hence, at some depth below the euphotic zone, usually around 1,000–1,500 m, the

ocean is generally rich in dissolved nutrients and a nutrient maximum is found

(Fig. 1.9b). This nutrient maximum corresponds to the midwater oxygen minimum

(Fig. 1.9a). The transition zone between the base of the surface layer where nutrient

concentrations are low and the depth of the nutrient maximum is called the nutricline. Similar changes in oxygen and nutrient concentrations occur as ocean

water circulates between the different ocean basins. Thus as deep water moves from

its formation region in the North Atlantic via the South Atlantic to the North Pacific

Ocean, the dissolved oxygen concentrations slowly decrease and the concentrations

of dissolved macro nutrients (N, P and Si) and carbon dioxide increase, so that the

deep North Pacific has the highest overall dissolved nutrient concentrations

(Fig. 1.9b).

Only a small fraction (<0.1 %) of sinking particulate matter reaches the seafloor

in the open ocean (Martin et al. 1987). Hence, a large fraction of nutrients is recycled

in the ocean interior. The physical process of upwelling (i.e., the upward movement

of water parcels) plays a fundamental role in marine ecosystems as it lifts

nutrient-enriched deeper water into the euphotic zone stimulating photosynthesis.



16



1 Preliminaries



Fig. 1.9 Typical vertical

distributions of dissolved

oxygen and nitrate nitrogen

(as an example of a key

nutrient) in different oceanic

regions. Adapted from Segar

(2007)



1.5.5



Nutrient Limitation



The limitation of phytoplankton growth has traditionally been interpreted in the

context of Liebig’s Law of the Minimum (Sprengel 1828; Liebig 1855), which states

that plant growth will be as great as allowed by the least available resource, the

limiting nutrient that sets the productivity of the system (de Baar 1994). In the

ocean, nitrogen, in the form of nitrate, is usually considered to limit production.

This is in contrast to freshwater systems, where phosphorous is usually considered

to be the limiting nutrient. However, recent studies in the eutrophic waters off the

Louisiana shelf have shown that phosphorous can be limiting here at certain times

(Sylvan et al. 2006; Quigg et al. 2011.)

In the 1930s the English biologist, Joseph Hart, speculated that the ocean’s great

“desolate zones” (areas apparently rich in nutrients, but lacking in plankton activity

or other sea life) might simply be iron deficient (Weier 2001). Little further scientific discussion of this issue was recorded until the 1980s, when oceanographer

John Martin renewed the controversy on the topic with his nutrient analyses of

seawater. His studies indicated it was indeed a scarcity of the micronutrient iron that

was limiting phytoplankton growth and overall productivity in these “desolate”

regions, which came to be called “High Nutrient, Low Chlorophyll” (HNLC) zones

(Martin and Fitzwater 1988; Boyd et al. 2007). These represent about 40–50 % of

the areal extent of the world’s oceans (Moore et al. 2002).

Iron limitation has been identified for the upwelling regions of the Humboldt

Current (Hutchins et al. 2002); and the California Current (Hutchins et al. 1998;

Chase et al. 2007). The role of external nutrient input, particularly iron, via

atmospheric dust plumes became apparent from observations in the Canary Current

upwelling system (Neuer et al. 2004). In upwelling regions, the continual importation of deeper waters, either along a “line feature” in areas where there is a belt of



1.5 Light, Nutrients and Oxygen in the Sea



17



upwelling (such as Oregon), or more usually at upwelling centres, as found for

example off California or in the Benguela region, means that this limitation is

generally removed, and indeed, silica has been suggested as limiting production in

the equatorial Pacific and off Peru (Dugdale et al. 1995). Silicon availability can

also be a major limiting factor in polar regions (e.g., Nelson and Tréguer 1992).

While this view of nutrient limitation is powerful, interactions among nutrients

and between nutrients and light can also control productivity. A simple but

important example of this potential for “co-limitation” comes from polar regions,

where oblique solar insolation combines with deep mixing of surface waters to

yield low light levels. In such environments, higher iron supply can increase the

efficiency with which phytoplankton capture light energy (Sunda and Huntsman

1997; Maldonado et al. 1999). More broadly, it has been argued that phytoplankton

generally reside in a state of co-limitation by all the chemicals they require,

including the many trace metal nutrients (Morel 2008).



1.5.6



Mechanisms Limiting Phytoplankton Blooms



As stated above, there are three key factors controlling primary production in the

surface ocean. The first factor is the maximum light intensity and, hence, euphotic

depth which underlies seasonal changes. Light intensity can be dramatically

reduced in coastal regions by continental influences (e.g., sediment resuspension or

sediment inputs from rivers).

The second factor is the depth of the surface mixed layer. Phytoplankton and

other organic matter are vertically stirred throughout the surface mixed layer.

Hence, the depth of surface mixed layer influences the relative time of growth that

such organisms have when moving through the euphotic zone. The overall depth of

the surface mixed layer depends on surface heat and freshwater fluxes and the

magnitude of the wind stress, which are seasonally variable. The mixed-layer depth

can vary on time scales from minutes to weeks under the influence of storm-induced

mixing, internal waves and upwelling processes.

The third factor is the concentration level of nutrients within the euphotic zone,

which depends on both physical processes (i.e., nutrient supply from mixed-layer

deepening or upwelling) and biological processes (i.e., nutrient consumption for

primary production). Nutrients become rapidly exhausted in the surface mixed layer

via consumption unless there is an external nutrient source. This includes the

upwelling process, which reduces the depth of the surface mixed layer and lifts

elevated nutrient levels closer to the sea surface, and mixed-layer deepening from

storms or thermohaline convection which entrains nutrient-enriched sub-pycnocline

water into the surface mixed layer. Hence, light intensity, nutrient distributions and

mixed layer depth all operate together to control primary production in the surface

ocean. Dramatic reduction in oxygen levels during excessive algal growth can be



18



1 Preliminaries



Fig. 1.10 Schematic of the

situation in which light and

nutrient limitations created a

zone of sub-surface

phytoplankton production

confined to the base of the

surface mixed layer



another controlling factor. Figure 1.10 shows the creation of a zone of sub-surface

phytoplankton production which is confined by too low nutrient concentrations

above and too low light intensity below.



1.5.7



Nutrient Regeneration



Benthic nutrient regeneration from ammonia was first explored and described by

Dugdale and Goering (1967) and later by Eppley (1992) for the California Current

upwelling system. This led to the introduction of the f ratio, the ratio between new

and regenerated production, by Eppley and Peterson (1979). The f ratio plays a key

role in the characterization of upwelling systems.

Nutrient regeneration plays a significant role in upwelling regions. It takes place

primarily through two processes, bacterial regeneration at the sediment-water interface and in the water column, and by grazing activities of herbivores (Fig. 1.11

illustrates the regeneration process). The supply of nitrogen as dissolved nitrate

allows us to differentiate between “new” production fueled by nitrate and “regenerated” production fueled by recycled ammonium and urea. In the open ocean, the

f-ratio is generally about 0.1. In coastal upwelling regions, however, it can be as high

as 0.8 (Laws 2004).

The increased productivity of upwelling regions results directly from the continuing availability of upwelled nitrate for new production, in contrast to other

coastal regions or the open ocean that rely on much smaller quantities of recycled

nitrogen. The fractions of regeneration attributable to herbivores and to bacterial

action vary from region to region (Dugdale 1972). Early analyses of Dugdale and

Goering (1970) show that regeneration of nitrogen and phosphorous by the

anchoveta populations in the Peru upwelling system take place at such high rates



1.5 Light, Nutrients and Oxygen in the Sea



19



Fig. 1.11 The approximate pathways of phosphorous, nitrogen, and silica circulation, and

biological uptake and regeneration in an upwelling region. Redrawn after Dugdale (1972)



that the anchoveta must be the dominant regenerators there. Direct silica regeneration was found to take place through anchoveta grazing activities at 10–20 % of

the rate for nitrogen.



1.6

1.6.1



The Carbon Cycle and Oceanic Carbon Pumps

Overview



The initial source of carbon on Earth is outgassing of CO2, stored in the mantle

when the Earth was formed, from the Earth’s interior at mid-ocean ridges or hotspot

volcanoes. A second source is found at subduction-related volcanic arcs, and most

CO2 released at these subduction zones is derived from the metamorphism of

sedimentary carbonate rocks subducting with the ocean crust. On geological

timescales (millions of years), carbon is released into the atmosphere and ocean

through the weathering of carbonate rocks such as limestone and via volcanic

emissions. It returns as new rocks formed through sediment deposition.

On the much shorter timescale <100 years, carbon is exchanged between the

atmosphere, the ocean and living and dead organisms, and air-sea gas exchange is the

major process controlling carbon-dioxide fluxes across the sea surface. From the start

of the industrial revolution in the mid-18th century, the atmospheric carbon budget

has been substantially disturbed through human activities, such as fossil fuel combustion and cement manufacture, so that the pre-industrial atmospheric CO2 concentration of about 270 ppm now exceeds 403 ppm and is continuing to increase.



20



1 Preliminaries



Roughly 50 % of the CO2 produced by human activities is taken up by the ocean, the

remainder staying in the atmosphere where it contributes to global warming.

Atmospheric CO2 enters the ocean via air-sea gas transfer. This transfer is a

function of a transfer coefficient, called piston velocity, and the difference in partial

gas pressures across the sea surface. Under the assumption that the thin surface skin

of the ocean is fully saturated with a gas and applying Henry’s and Fick’s laws, the

air-sea gas flux can be formulated as:





F ẳ u kH Pgas Cml ;



1:3ị



where u is the transfer coefficient, which is strongly controlled by wind speed, kH is

solubility, Pgas is partial pressure of the gas in the atmosphere, and Cml is the gas

concentration in the surface mixed layer of the ocean. Equation (1.3) is known as

the film model of gas exchange. The direction of the gas flux depends on whether

the dissolved gas in the mixed layer is under-saturated (F > 0), when CO2 will enter

the ocean from the atmosphere, or oversaturated (F < 0), when the gas will move in

the opposite direction. As for oxygen, the solubility of carbon dioxide decreases

with increasing sea temperature. Its saturation concentration value more than

doubles as temperatures decrease from 24 °C (tropical regions) to 0 °C (polar

regions). Hence, fully-saturated cold water can hold more CO2 than warm water

and the deep oceans have higher concentrations of CO2 than the surface layers.

Within the ocean, the general carbon cycle is a complex process driven by both

biogeochemistry and physics (Fig. 1.12). The oceanic carbon cycle can be

described by different carbon pumps, each describing specific mechanisms that

transfer carbon dioxide from the upper to the deep ocean or vice versa. These

pumps are the solubility pump and the biological pump. The solubility pump is

responsible for about 20 % of the vertical gradient in dissolved inorganic carbon in

the ocean, while the remaining 80 % originates from the biological pump

(Sarmiento et al. 1995). The solubility pump operates as a combination of:

(a) the temperature dependency of the solubility of carbon dioxide in seawater;

i.e. under the same atmospheric conditions, cold water can dissolve more CO2 than

warm water before it reaches an equilibrium with the atmosphere, and

(b) oceanic flows that either export surface water to the ocean interior (called

oceanic subduction) or bring deeper water back to the sea surface (upwelling).

One branch of the solubility pump, for example, is the deep circulation of the

oceans driven by open-ocean convection in sub-polar regions of the North Atlantic

Ocean. The other branch is the reverse process of upwelling in which CO2 enriched

deeper water is returned to the sea surface.

The biologic pump (responsible for 80 % of total carbon fixation in the ocean)

describes vertical carbon transfers in the ocean associated with biochemical processes. The biologic pump comprises:



1.6 The Carbon Cycle and Oceanic Carbon Pumps



21



Fig. 1.12 The basic oceanic carbon cycle. Adapted from Lalli and Parsons (1993)



(a) the organic carbon pump, associated with primary production in the euphotic

zone and remineralization of detritus at depths, and

(b) the calcium carbonate counter pump, associated with skeleton and shell

formation in the surface ocean and the dissolution of calcareous particles at depth.

The biological pump starts with the conversion of inorganic carbon to organic

forms. Some of the phytoplankton are remineralized when they die, but the major

fraction is consumed by zooplankton and nekton, some of which also take up carbon

dioxide directly to form calcium carbonate shells. Zooplankton faecal material and

dead phytoplankton cells sink, transferring carbon into the deeper ocean, and remineralization continues throughout the water column. While a small portion (<0.1 %)

of the carbon can eventually be preserved in ocean sediments, most is remineralized

into carbon dioxide below 300 m depth, after which upwelling and the general

circulation eventually brings it back to the surface layer as bicarbonate or carbonate

ions, from where some returns to the atmosphere as carbon dioxide gas. Although

only a small percentage in terms of the total mass of carbon at a given instant, over

geological time the preservation of the carbonate skeletons of marine organisms is an

extremely important component of the global carbon cycle, with about 1,000 times

as much carbon sequestered in limestone or organic marine sediments as exists as

free CO2, bicarbonate or carbonate ions (Lalli and Parsons 1993).



22



1 Preliminaries



When carbon dioxide from the atmosphere reacts with seawater (H2O), it

immediately forms a weak acid, carbonic acid (H2CO3), which in itself is chemically unstable. This acid further dissociates to form bicarbonate HCO3− (a base) and

hydrogen ions H+ (an acid):

CO2 ỵ H2 O ! H2 CO3 ! HCO3 ỵ H ỵ



1:4ị



Excess hydrogen ions (H+) react with carbonate ions (CO32−) (seawater is naturally saturated with this base) to form further bicarbonate ions:

H ỵ ỵ CO3 2 ! 2HCO3 À



ð1:5Þ



The acidity of the oceans is determined by the concentration of hydrogen ions; a

greater amount results in more acidic conditions, represented by a lower pH.

The carbonate ions (CO32−) and bicarbonate ions (HCO3−) can react with calcium ions, which are in excess in seawater, to form calcium carbonate (CaCO3)

which underpins skeleton and shell formation (also known as calcification) in

marine organisms such as corals, shellfish and marine plankton (Feely et al. 2008).

The main chemical reactions for the mineral formation and the dissolution of calcium carbonate (CaCO3) are as follows:

CaCO3 $ Ca2 þ þ CO3 2À



ð1:6Þ



CaCO3 þ H2 O þ CO2 $ Ca2 ỵ ỵ 2HCO3



1:7ị



Calcium carbonate is formed as the reaction proceeds from right to left, and

dissolved from left to right. In contrast to the organic carbon pump, the calcification

process releases CO2 back into the ambient seawater and dissolution of calcareous

particles at depth takes up dissolved CO2. The reaction of CO2 with seawater to

form bicarbonate and carbonate ions means that the resultant increase in gaseous

seawater CO2 concentration is smaller than the actual amount of carbon dioxide

entering the seawater. This chemical reaction together with the gravitational export

of detritus from the euphotic zone supports a continuous air-sea gas transfer of CO2

into the ocean and is quantitatively the most important oceanic process contributing

to the ocean as an overall carbon sink (Feely et al. 2008). The buffering capacity of

seawater also implies that seawater maintains a slightly basic pH state within relatively narrow limits, despite the uptake of atmospheric CO2, although this appears

to be changing towards less basic conditions.

The calcification process depends critically on the availability of two specific

carbonate minerals, aragonite and calcite. Aragonite is used by pteropods to construct their shells, while calcite is used by coccoliths and foraminifera. When

seawater is supersaturated with these minerals, as is the case in all ocean surface

waters at present, the formation of shells and skeletons will be favoured.

Conversely, when seawater is under-saturated with respect to these minerals, the

seawater becomes corrosive and the shells of calcifying organisms are increasingly

prone to dissolution (Feely et al. 1988).



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