What is Ocean Productivity?

Ocean productivity largely refers to the production of organic matter by "phytoplankton," plants suspended in the ocean, most of which are single-celled. Phytoplankton are "photoautotrophs," harvesting light to convert inorganic to organic carbon, and they supply this organic carbon to diverse "heterotrophs," organisms that obtain their energy solely from the respiration of organic matter. Open ocean heterotrophs include bacteria as well as more complex single- and multi-celled "zooplankton" (floating animals), "nekton" (swimming organisms, including fish and marine mammals), and the "benthos" (the seafloor community of organisms).

The many nested cycles of carbon associated with ocean productivity are revealed by the following definitions (Bender et al. 1987) (Figure 1). "Gross primary production" (GPP) refers to the total rate of organic carbon production by autotrophs, while "respiration" refers to the energy-yielding oxidation of organic carbon back to carbon dioxide. "Net primary production" (NPP) is GPP minus the autotrophs' own rate of respiration; it is thus the rate at which the full metabolism of phytoplankton produces biomass. "Secondary production" (SP) typically refers to the growth rate of heterotrophic biomass. Only a small fraction of the organic matter ingested by heterotrophic organisms is used to grow, the majority being respired back to dissolved inorganic carbon and nutrients that can be reused by autotrophs. Therefore, SP in the ocean is small in comparison to NPP. Fisheries rely on SP; thus they depend on both NPP and the efficiency with which organic matter is transferred up the foodweb (i.e., the SP/NPP ratio). "Net ecosystem production" (NEP) is GPP minus the respiration by all organisms in the ecosystem. The value of NEP depends on the boundaries defined for the ecosystem. If one considers the sunlit surface ocean down to the 1% light level (the "euphotic zone") over the course of an entire year, then NEP is equivalent to the particulate organic carbon sinking into the dark ocean interior plus the dissolved organic carbon being circulated out of the euphotic zone. In this case, NEP is also often referred to as "export production" (or "new production" (Dugdale & Goering 1967), as discussed below). In contrast, the NEP for the entire ocean, including its shallow sediments, is roughly equivalent to the slow burial of organic matter in the sediments minus the rate of organic matter entering from the continents.

Figure 1

Productivity in the surface ocean, the definitions used to describe it, and its connections to nutrient cycling. The blue cycle for “net ecosystem production” (NEP) (i.e. “new” or “export” production) encompasses the “new” nutrient supply from the ocean interior, its uptake by autotrophic phytoplankton growth, packaging into large particles by heterotrophic grazing organisms, and sinking of organic matter out of the surface ocean. The red cycle illustrates the fate of the majority of organic matter produced in the surface ocean, which is to be respired by heterotrophic organisms to meet their energy requirements, thereby releasing the nutrients back into the surface water where they can be taken up by phytoplankton once again to fuel “regenerated production.” The green cycle represents the internal respiration of phytoplankton themselves, that is, their own use of the products of photosynthesis for purposes other than growth. These nested cycles combine to yield (1) “gross primary production” (GPP) representing the gross photosynthesis and (2) “net primary production” (NPP) that represents phytoplankton biomass production that forms the basis of the food web plus a much smaller rate of organic matter export from the surface. While the new nutrient supply and export production are ultimately linked by mass balance, there may be imbalances on small scales of space and time, allowing for brief accumulations of biomass.

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There are no accumulations of living biomass in the marine environment that compare with the forests and grasslands on land (Sarmiento & Bender 1994). Nevertheless, ocean biology is responsible for the storage of more carbon away from the atmosphere than is the terrestrial biosphere (Broecker 1982). This is achieved by the sinking of organic matter out of the surface ocean and into the ocean interior before it is returned to dissolved inorganic carbon and dissolved nutrients by bacterial decomposition. Oceanographers often refer to this process as the "biological pump," as it pumps carbon dioxide (CO₂) out of the surface ocean and atmosphere and into the voluminous deep ocean (Volk & Hoffert 1985).

Only a fraction of the organic matter produced in the surface ocean has the fate of being exported to the deep ocean. Of the organic matter produced by phytoplankton (NPP), most is respired back to dissolved inorganic forms within the surface ocean and thus recycled for use by phytoplankton (Eppley & Peterson 1979) (Figure 1). Most phytoplankton cells are too small to sink individually, so sinking occurs only once they aggregate into larger particles or are packaged into "fecal pellets" by zooplankton. The remains of zooplankton are also adequately large to sink. While sinking is a relatively rare fate for any given particle in the surface ocean, biomass and organic matter do not accumulate in the surface ocean, so export of organic matter by sinking is the ultimate fate for all of the nutrients that enter into the surface ocean in dissolved form — with the exceptions that (1) dissolved nutrients can be returned unused to the interior by the circulation in some polar regions (see below), and (2) circulation also carries dissolved organic matter from the surface ocean into the interior, a significant process (Hansell et al. 2009) that we will not address further. As organic matter settles through the ocean interior and onto the seafloor, it is nearly entirely decomposed back to dissolved chemicals (Emerson & Hedges 2003, Martin et al. 1987). This high efficiency of decomposition is due to the fact that the organisms carrying out the decomposition rely upon it as their sole source of chemical energy; in most of the open ocean, the heterotrophs only leave behind the organic matter that is too chemically resistant for it to be worth the investment to decompose. On the whole, only a tiny fraction (typically much less than 1%) of the organic carbon from NPP in the euphotic zone survives to be buried in deep sea sediments.

Productivity in coastal ecosystems is often distinct from that of the open ocean. Along the coasts, the seafloor is shallow, and sunlight can sometimes penetrate all the way through the water column to the bottom, thus enabling bottom-dwelling ("benthic") organisms to photosynthesize. Furthermore, sinking organic matter isintercepted by the seabed, where it supports thriving benthic faunal communities, in the process being recycled back to dissolved nutrients that are then immediately available for primary production. The proximity to land and its nutrient sources, the interception of sinking organic matter by the shallow seafloor, and the propensity for coastal upwelling all result in highly productive ecosystems. Here, we mainly address the productivity of the vast open ocean; nevertheless, many of the same concepts, albeit in modified form, apply to coastal systems.

What Does Ocean Productivity Need?

Phytoplankton require a suite of chemicals, and those with the potential to be scarce in surface waters are typically identified as "nutrients." Calcium is an example of an element that is rapidly assimilated by some plankton (for production of calcium carbonate "hard parts") but is not typically considered a nutrient because of its uniformly high concentration in seawater. Dissolved inorganic carbon, which is the feedstock for organic carbon production by photosynthesis, is also abundant and so is not typically listed among the nutrients. However, its acidic form dissolved CO₂ is often at adequately low concentrations to affect the growth of at least some phytoplankton.

Broadly important nutrients include nitrogen (N), phosphorus (P), iron (Fe), and silicon (Si). There appear to be relatively uniform requirements for N and P among phytoplankton. In the early 1900s, oceanographer Alfred Redfield found that plankton build their biomass with C:N:P stoichiometric ratios of ~106:16:1, to which we now refer as the Redfield ratios (Redfield 1958). As Redfield noted, the dissolved N:P in the deep ocean is close to the 16:1 ratio of plankton biomass, and we will argue below that plankton impose this ratio on the deep, not vice versa. Iron is found in biomass only in trace amounts, but it is used for diverse essential purposes in organisms, and it has become clear over the last 25 years that iron's scarcity often limits or affects productivity in the open ocean, especially those regions where high-N and -P deep water is brought rapidly to the surface (Martin & Fitzwater 1988). Research is ongoing to understand the role of other trace elements in productivity (Morel et al. 2003). Silicon is a nutrient only for specific plankton taxa-diatoms (autotrophic phytoplankton), silicoflaggellates, and radiolaria (heterotrophic zooplankton) — which use it to make opal hard parts. However, the typical dominance of diatoms in Si-bearing waters, and the tendency of diatom-associated organic matter to sink out of the surface ocean, make Si availability a major factor in the broader ecology and biogeochemistry of surface waters.

Sunlight is the ultimate energy source — directly or indirectly — for almost all life on Earth, including in the deep ocean. However, light is absorbed and scattered such that very little of it penetrates below a depth of ~80 m (as deep as 150 m in the least productive subtropical regions, but as shallow as 10 m in highly productive and coastal regions) (Figure 2). Thus, photosynthesis is largely restricted to the upper light-penetrated skin of the ocean. Moreover, across most of the ocean's area, including the tropics, subtropics, and the temperate zone, the absorption of sunlight causes surface water to be much warmer than the underlying deep ocean, the latter being filled with water that sank from the surface in the high latitudes . Warm water is more buoyant than cold, which causes the upper sunlit layer to float on the denser deep ocean, with the transition between the two known as the "pycnocline" (for "density gradient") or "thermocline" (the vertical temperature gradient that drives density stratification across most of the ocean, Figure 2). Wind or another source of energy is required to drive mixing across the pycnocline, and so the transport of water with its dissolved chemicals between the sunlit surface and the dark interior is sluggish. This dual effect of light on photosynthesis and seawater buoyancy is critical for the success of ocean phytoplankton. If the ocean did not have a thin buoyant surface layer, mixing would carry algae out of the light and thus away from their energy source for most of the time. Instead of nearly neutrally buoyant single celled algae, larger, positively buoyant photosynthetic organisms (e.g., pelagic seaweeds) might dominate the open ocean. This hypothetical case aside, although viable phytoplankton cells are found (albeit at low concentrations) in deeper waters, photosynthesis limits active phytoplankton growth to the upper skin of the ocean, while upper ocean density stratification prevents them from being mixed down into the dark abyss. Thus, most open ocean biomass, including phytoplankton, zooplankton, and nekton, is found within ~200 m of the ocean surface.

Figure 2

Typical conditions in the subtropical ocean, as indicated by data collected at the Bermuda Atlantic Time-series Station in July, 2008. The thermocline (vertical temperature gradient) stratifies the upper water column. During this particular station occupation, the shallow wind-mixed surface layer is not well defined, presumably because of strong insolation and a lack of wind that allowed continuous stratification all the way to the surface. Very little sunlight penetrates deeper than ~100 m. New supply of the major nutrients N and P is limited by the slow mixing across the upper thermocline (showing here only the N nutrient nitrate, NO₃^-). Within the upper euphotic zone, the slow nutrient supply is completely consumed by phytoplankton in their growth. This growth leads to the accumulation of particulate organic carbon in the surface ocean, some of which is respired by bacteria, zooplankton, and other heterotrophs, and some of which is exported as sinking material. The deep chlorophyll maximum (DCM) occurs at the contact where there is adequate light for photosynthesis and yet significant nutrient supply from below. The DCM should not be strictly interpreted as a depth maximum in phytoplankton biomass, as the phytoplankton at the DCM have a particularly high internal chlorophyll concentration. The data shown here is made available the Bermuda Institute of Ocean Sciences (http://bats.bios.edu) and the Bermuda Bio Optics Project (http://www.icess.ucsb.edu/bbop/).

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At the same time, the existence of a thin buoyant surface layer conspires with other processes to impose nutrient limitation on ocean productivity. The export of organic matter to depth depletes the surface ocean of nutrients, causing the nutrients to accumulate in deep waters where there is no light available for photosynthesis (Figure 2). Because of the density difference between surface water and the deep sea across most of the ocean, ocean circulation can only very slowly reintroduce dissolved nutrients to the euphotic zone. By driving nutrients out of the sunlit, buoyant surface waters, ocean productivity effectively limits itself.

Phytoplankton growth limitation has traditionally been interpreted in the context of Liebig's Law of the Minimum, 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). While this view 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 availability. In such environments, higher iron supply can increase the efficiency with which phytoplankton capture light energy (Maldonado et al. 1999, Sunda & Huntsman 1997). More broadly, it has been argued that phytoplankton should generally seek a state of co-limitation by all the chemicals they require, including the many trace metal nutrients (Morel 2008).

Who Are the Major Players in Ocean Productivity?

In contrast to the terrestrial biosphere, most marine photosynthesis is conducted by single-celled organisms, and the more abundant of the multicellular forms are structurally much simpler than the vascular plants on land. During much of the twentieth century, it was thought that cells in the range of ~5 to ~100 microns diameter account for most phytoplankton biomass and productivity. This size range is composed mostly of eukaryotes, organisms whose cells contain complex membrane-bound structures ("organelles"), including the cell's nucleus and chloroplasts. Well-studied forms of eukaryotic phytoplankton include the opal-secreting diatoms, prymnesiophytes (including the CaCO₃-secreting coccolithophorids), and the organic wall-forming dinoflagellates. The centrality of these organisms in early oceanographic thought was due to their accessibility by standard light microscopy.

Only with recent technological advances have smaller organisms become readily observable, revolutionizing our view of the plankton. In particular, the cyanobacteria, which are prokaryotes (lacking a nucleus and most other organelles found in eukaryotes), are now known to be important among the phytoplankton. Initially, the cyanobacteria were identified largely with colonial forms such as Trichodesmium that play the critical role of "fixing" nitrogen (see below). However, major discoveries over the last thirty years have revealed the prevalence across the global ocean of unicellular cyanobacteria of ~0.5 to ~1.5 microns diameter. It is now recognized that two cyanobacterial genera — Synechoccocus and Prochlorococcus — dominate phytoplankton numbers and biomass in the nutrient-poor tropical and subtropical ocean (Waterbury et al. 1979, Chisholm et al. 1988). In addition, new methods, both microscopic and genetic, are revealing a previously unappreciated diversity of smaller eukaryotes in the open ocean.

Mapping ecological and biogeochemical functions onto the genetic diversity of the phytoplankton is an active area in biological and chemical oceanography. Based on observations as well as theory, the smaller phytoplankton such as the unicellular cyanobacteria are thought to dominate regenerated production in many systems, whereas the larger eukaryotes appear to play a more important role in new production (i.e., NEP, Figure 1; see below).

The food source of a given form of zooplankton is typically driven by its own size, with microzooplankton grazing on the prokaryotes and smaller eukaryotes and multicellular zooplankton grazing on larger eukaryotes, both phytoplankton and microzooplankton. Because of their relative physiological simplicity, microzooplankton are thought to be highly efficient grazers that strongly limit the biomass accumulation of their prey. In contrast, the multicellular zooplankton, because they typically have more complex life histories, can lag behind the proliferation of their prey, allowing them to bloom and sometimes avoid predation altogether and sink directly. The multicellular zooplankton also often facilitate the production of sinking organic matter, for example, through the production of fecal pellets by copepods.

The diversity of the plankton interacts with open ocean environmental conditions to affect the productivity of the larger ecosystem (Michaels & Silver 1988, Morel et al. 1991, Buesseler 1998) (Figure 3). In the nutrient-poor tropical and subtropical ocean, the (small) cyanobacteria tend to be numerically dominant, perhaps because they specialize in taking up nutrients at low concentrations. Small phytoplankton have a greater surface area-to-volume ratio than do large phytoplankton. A greater proportional surface area promotes the uptake of nutrients across the cell boundary, a critical process when nutrients are scarce, likely explaining why small phytoplankton dominate the biomass in the nutrient-poor ocean. The microzooplankton effectively graze these small cells, preventing their biomass from accumulating and sinking directly. Moreover, these single-celled microzooplankton lack a digestive tract, so they do not produce the fecal pellets that represent a major mechanism of export. Instead, any residual organic matter remains in the upper ocean, to be degraded by bacteria. All told, microzooplankton grazing of phytoplankton biomass leads to the remineralization of most of its contained nutrients and carbon in the surface ocean, and thus increases recycling relative to organic matter export. This very efficient recycling elevates NPP relative to NEP, yielding a low NEP:NPP ratio (~0.05–0.3) in nutrient-poor systems (Figure 3a). In contrast, larger phytoplankton, such as diatoms, often dominate the nutrient-rich polar ocean, and these can be grazed directly by multicellular zooplankton. By growing adequately rapidly to outstrip the grazing rates of these zooplankton, the diatoms can sometimes accumulate to high concentrations and produce abundant sinking material. In addition, the zooplankton export organic matter as fecal pellets. In these productive systems, the less intensive upper ocean recycling causes NEP and NPP to be more similar, with an NEP:NPP ratio often near 0.5 (Figure 3b).

Figure 3

The most broadly accepted paradigm for the controls on surface nutrient recycling efficiency. NPP is supported by both new nutrient supply from the deep ocean and nutrients regenerated within the surface ocean. The fraction of NEP:NPP ratio appears to vary with the nutrient supply, because links to the ecology of the plankton. In the nutrient-poor tropical and subtropical ocean (a), the (small) cyanobacteria tend to be numerically dominant. The microzooplankton that graze these small cells do so effectively, preventing phytoplankton from sinking directly. Moreover, these single-celled microzooplankton do not produce sinking fecal pellets. Instead, any residual organic matter remains to be degraded by bacteria. This increases recycling relative to organic matter export, yielding a low NEP:NPP ratio (~0.1). In nutrient-rich regions (b), large phytoplankton are more important, and these can be grazed directly by multicellular zooplankton. By growing adequately rapidly to outstrip the grazing rates of zooplankton, the large phytoplankton can sometimes accumulate to high concentrations and produce abundant sinking material. In addition, the zooplankton export organic matter as fecal pellets. In these productive systems, the less intensive upper ocean recycling causes NEP and NPP to be more similar, with an NEP:NPP ratio often near 0.5. The relationships between nutrient supply, phytoplankton size, and sinking thus dominate this view of upper ocean nutrient cycling.

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How Does Ocean Productivity Vary in the Modern Ocean?

Satellites can measure the color of the surface ocean in order to track the concentration of the green pigment chlorophyll that is used to harvest light in photosynthesis (Figure 4). Higher chlorophyll concentrations and in general higher productivity are observed on the equator, along the coasts (especially eastern margins), and in the high latitude ocean (Figure 4a and b). A major driver of these patterns is the upwelling and/or mixing of high nutrient subsurface water into the euphotic zone, as is evident from surface nutrient measurements (Figure 4c and d).

Figure 4

Composite global ocean maps of concentrations of satellite-derived chlorophyll and ship-sampled nitrate (NO₃^-; the dominant N-containing nutrient). Northern hemisphere summer is shown in the left panels and southern hemisphere summer on the right. In the vast unproductive low- and mid-latitude ocean, warm and sunlit surface water is separated from cold, nutrient-rich interior water by a strong density difference that restricts mixing of water and thereby reduces nutrient supply, which becomes the limiting factor for productivity. These "ocean deserts" are dissected by areas, mainly at the equator and the eastern margins of ocean basins, where the wind pushes aside the buoyant, warm surface lid and allows nutrient-rich deeper water to be upwelled. In the high latitude ocean, surface water is cold and therefore the vertical density gradient is weak, which allows for vertical mixing of water to depths much greater than the sunlit "euphotic zone" as a result, the nutrient supply is greater than the phytoplankton can consume, given the available light (and iron, see text). The data shown here are available through the NASA’s OceanColor (http://oceancolor.gsfc.nasa.gov/) and NOAA’s National Oceanographic Data Center (http://www.nodc.noaa.gov/) websites. Sea ice cover impedes measurement of ocean color from space, reducing the apparent areas of the polar oceans in the winter hemisphere (upper panels).

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There are caveats regarding the use of satellite-derived chlorophyll maps to deduce productivity, phytoplankton abundance, and their variation. First, the relationship between chlorophyll and biomass is changeable, depending on the physiology of phytoplankton; for example, phytoplankton adapted to lower light and/or higher nutrients (e.g., iron) tend to have a higher cellular concentration of chlorophyll (Geider et al. 1997). Second, chlorophyll concentration speaks more directly to the rate of photosynthesis (i.e., GPP) than to NPP, the latter representing the growth of phytoplankton biomass plus the transfer of organic matter-bound energy to higher trophic levels. Third, for a given NPP, small variations in grazing can lead to large proportional changes in phytoplankton biomass (Landry & Hassett 1982). Fourth, the depth range sensed by the satellite ocean color measurements extends only to the uppermost ten's of meters, much shallower than the base of the euphotic zone (Figure 2). Compared to nutrient-bearing regions, nutrient-deplete regions (e.g., the subtropical gyres) have a larger fraction of chlorophyll below the depth that can be sensed by the satellite (Smith 1981). Thus, satellite chlorophyll observations tend to over-accentuate the productivity differences between nutrient-bearing and -depleted regions. Despite these caveats, satellite-derived ocean color observations have transformed our view of ocean productivity.

In some temperate and subpolar regions, productivity reaches a maximum during the spring as the phytoplankton transition from light to nutrient limitation. In the highest latitude settings, while the "major nutrients" N and P remain at substantial concentrations, the trace metal iron can become limiting into the summer (Boyd et al. 2007, Martin & Fitzwater 1988). In at least some of these polar systems, it appears that light and iron can "co-limit" summertime photosynthesis (Maldonado et al. 1999, Mitchell et al. 1991).

In the Following Section: What Controls Ocean Productivity on Long Time Scales?