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5 Light, Nutrients and Oxygen in the Sea
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;
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
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
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)
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
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).
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
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.
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
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
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).
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
Fig. 1.10 Schematic of the
situation in which light and
nutrient limitations created a
zone of sub-surface
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.
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
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.
The Carbon Cycle and Oceanic Carbon Pumps
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.
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 ;
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
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).
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
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
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
CaCO3 þ H2 O þ CO2 $ Ca2 ỵ ỵ 2HCO3
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).