8.05 Biogeochemistry of Primary Production in the Sea Biogeochemistry of Primary Production in the Sea P. G. Falkowski Rutgers University, New Brunswick, NJ, USA 8.05.1 INTRODUCTION

8.05Biogeochemistry of PrimaryProduction in the SeaP. G. Falkowski

Rutgers University, New Brunswick, NJ, USA

8.05.1 INTRODUCTION 1868.05.1.1 The Two Carbon Cycles 1868.05.1.2 A Primer on Redox Chemistry 186

8.05.2 CHEMOAUTOTROPHY 187

8.05.3 PHOTOAUTOTROPHY 1888.05.3.1 Selective Forces in the Evolution of Photoautotrophy 1888.05.3.2 Selective Pressure in the Evolution of Oxygenic Photosynthesis 189

8.05.4 PRIMARY PRODUCTIVITY BY PHOTOAUTOTROPHS 1898.05.4.1 What are Photoautotrophs? 189

8.05.4.1.1 The red and green lineages 1918.05.4.2 Estimating Chlorophyll Biomass 191

8.05.4.2.1 Satellite based algorithms for ocean color retrievals 1938.05.4.3 Estimating Net Primary Production 194

8.05.4.3.1 Global models of net primary production for the ocean 1948.05.4.4 Quantum Efficiency of NPP 195

8.05.4.4.1 Comparing efficiencies for oceanic and terrestrial primary production 196

8.05.5 EXPORT, NEW AND “TRUE NEW” PRODUCTION 1968.05.5.1 Steady-state versus Transient State 197

8.05.6 NUTRIENT FLUXES 1988.05.6.1 The Redfield Ratio 198

8.05.7 NITRIFICATION 1988.05.7.1 Carbon Burial 1998.05.7.2 Carbon Isotope Fractionation in Organic Matter and Carbonates 1998.05.7.3 Balance between Net Primary Production and Losses 1998.05.7.4 Carbon Burial in the Contemporary Ocean 2018.05.7.5 Carbon Burial in the Precambrian Ocean 201

8.05.8 LIMITING MACRONUTRIENTS 2018.05.8.1 The Two Concepts of Limitation 201

8.05.9 THE EVOLUTION OF THE NITROGEN CYCLE 202

8.05.10 FUNCTIONAL GROUPS 2038.05.10.1 Siliceous Organisms 2048.05.10.2 Calcium Carbonate Precipitation 2048.05.10.3 Vacuoles 205

8.05.11 HIGH-NUTRIENT, LOW-CHLOROPHYLL REGIONS—IRON LIMITATION 206

8.05.12 GLACIAL–INTERGLACIAL CHANGES IN THE BIOLOGICAL CO2 PUMP 207

8.05.13 IRON STIMULATION OF NUTRIENT UTILIZATION 207

8.05.14 LINKING IRON TO N2 FIXATION 207

8.05.15 OTHER TRACE-ELEMENT CONTROLS ON NPP 208

8.05.16 CONCLUDING REMARKS 209

ACKNOWLEDGMENTS 209

REFERENCES 210

185

As the present condition of nations is the result ofmany antecedent changes, some extremely remoteand others recent, some gradual, others sudden andviolent, so the state of the natural world is the result ofa long succession of events, and if we would enlargeour experience of the present economy of nature, wemust investigate the effects of her operations informer epochs.

Charles Lyell,Principles of Geology, 1830

8.05.1 INTRODUCTION

Earth is the only planet in our solar system thatcontains vast amounts of liquid water on itssurface and high concentrations of free molecularoxygen in its atmosphere. These two features arenot coincidental. All of the original oxygen onEarth arose from the photobiologically catalyzedsplitting of water by unicellular photosyntheticorganisms that have inhabited the oceans for atleast 3 Gyr. Over that period, theseorganisms have used the hydrogen atoms fromwater and other substrates to form organic matterfrom CO2 and its hydrated equivalents. Thisprocess, the de novo formation of organic matterfrom inorganic carbon, or primary production, isthe basis for all life on Earth. In thischapter, we examine the evolution and biogeo-chemical consequences of primary production inthe sea and its relationship to other biogeochem-ical cycles on Earth.

8.05.1.1 The Two Carbon Cycles

There are two major carbon cycles on Earth.The two cycles operate in parallel. One cycle isslow and abiotic. Its effects are observed onmultimillion-year timescales and are dictated bytectonics and weathering (Berner, 1990). In thiscycle, CO2 is released from the mantle to theatmosphere and oceans via vulcanism and seafloorspreading, and removed from the atmosphere andocean primarily by reaction with silicates to formcarbonates in the latter reservoir. Most of thecarbonates are subsequently subducted into themantle, where they are heated, and their carbon isreleased as CO2 to the atmosphere and ocean, tocarry out the cycle again. The chemistry of thiscycle is dependent on acid–base reactions, andwould operate whether or not there was life on theplanet (Kasting et al., 1988). This slow carboncycle is a critical determinate of the concentrationof CO2 in Earth’s atmosphere and oceans ontimescales of tens and hundreds of millions ofyears (Kasting, 1993).

The second carbon cycle is dependent on thebiologically catalyzed reduction of inorganic

carbon to form organic matter, the overwhelmingmajority of which is oxidized back to inorganiccarbon by respiratory metabolism (Schlesinger,1997). This cycle, which is observable on time-scales of days to millenia, is driven by reduction—oxidation (redox) reactions that evolved over,2 Gyr, first in microbes, and subsequently inmulticellular organisms (Falkowski et al., 1998).A very small fraction of the reduced carbonescapes respiration and becomes incorporatedinto the lithosphere. In so doing, some of theorganic matter is transferred to the slow carboncycle. In this chapter, we will focus primarilyon this fast, biologically mediated carbon cycle inthe sea, and the supporting biogeochemicalprocesses and feedbacks.

8.05.1.2 A Primer on Redox Chemistry

The biologically mediated redox reactionscycle carbon through three mobile pools: theatmosphere, the ocean, and the biosphere. Ofthese, the ocean is by far the largest (Table 1);however, more than 98% of this carbon is foundin its oxidized state as CO2 and its hydratedequivalents, HCO3

2 and CO322. To form organic

molecules, the inorganic carbon must be chemi-cally reduced, a process that requires the additionof hydrogen atoms (not just protons, but protonsplus electrons) to the carbon atoms. Broadlyspeaking, these biologically catalyzed reductionreactions are carried out by two groups oforganisms, chemoautotrophs and photoauto-trophs, which are collectively called primary

Table 1 Carbon pools in the major reservoirs on Earth.

Pools Quantity(£1015 g)

Atmosphere 720

Oceans 38,400Total inorganic 37,400Surface layer 670Deep layer 36,730Dissolved organic 600

LithosphereSedimentary carbonates .60,000,000Kerogens 15,000,000

Terrestrial biosphere (total) 2,000Living biomass 600–1,000Dead biomass 1,200

Aquatic biosphere 1–2

Fossil fuels 4,130Coal 3,510Oil 230Gas 140Other (peat) 250

Biogeochemistry of Primary Production in the Sea186

producers. The organic carbon they synthesizefuels the growth and respiratory demands of theprimary producers themselves and all remainingorganisms in the ecosystem.

All redox reactions are coupled sequences.Reduction is accomplished by the addition of anelectron or hydrogen atom to an atom or molecule.In the process of donating an electron to anacceptor, the donor molecule is oxidized. Hence,redox reactions require pairs of substrates, andcan be described by a pair of partial reactions, orhalf-cells:

Aox þ nðe2Þ $ Are ð1aÞ

Bred 2 nðe2Þ $ Box ð1bÞ

The tendency for a molecule to accept or releasean electron is therefore “relative” to some othermolecule being capable of conversely releasingor binding an electron. Chemists scale thistendency, called the redox potential, E; relativeto the reaction

H2 $ 2Hþ þ 2e2 ð2Þ

which is arbitrarily assigned an E of 0 at pH 0, andis designated E0: Biologists define the redoxpotential at pH 7, 298 K (i.e., room temperature)and 1 atm pressure (¼101.3 kPa). When sodefined, the redox potential is denoted by thesymbols E0

0 or sometimes Em7: The E0

0 for astandard hydrogen electrode is 2420 mV.

8.05.2 CHEMOAUTOTROPHY

Organisms capable of reducing sufficient inor-ganic carbon to grow and reproduce in the darkwithout an external organic carbon source arecalled chemoautotrophs (literally, “chemical self-feeders”). Genetic analyses suggest that chemo-autotrophy evolved very early in Earth’s history,and is carried out exclusively by prokaryoticorganisms in both the Archea and Bacteriasuperkingdoms (Figure 1).

Early in Earth’s history, the biologicalreduction of inorganic carbon may have beendirectly coupled to the oxidation of H2. At present,however, free H2 is scarce on the planet’s surface.Rather, most of the hydrogen on the surface ofEarth is combined with other atoms, such as sulfuror oxygen. Activation energy is required to breakthese bonds in order to extract the hydrogen. Onesource of energy is chemical bond energy itself.For example, the ventilation of reduced mantlegases along tectonic plate subduction zones onthe seafloor provides hydrogen in the form ofH2S. Several types of microbes can couple theoxidation of H2S to the reduction of inorganiccarbon, thereby forming organic matter in theabsence of light.

Ultimately all chemoautotrophs depend on anonequilibrium redox gradient, without whichthere is no thermodynamic driver for carbonfixation. For example, the reaction involving theoxidation of H2S by microbes in deep-sea ventsdescribed above is ultimately coupled to oxygen

Bacteria

Proteobacteria

Cyanobacteria

Flamobacteria

Gram-positive bacteriaGreen non-sulfur bacteria

Archea

Three domains of lifeEukarya

AnimalsSlime Molds

Entamoebae

Methanobacterium

MethanococcusThermoproteus Thermococcus

Thermotogz

Aquifex

UniversalAncestor

DiplomanadsMicrosporidia

Euglena

Ciliates

Plants

Fungi

Pyrodictium

SuiflobusMethanopyrus

Archeo globus

Thermoplasma

Halobacterium

Pathways of autotrophic CO2 fixation

Reductive citric acid cycle

Reductive acetyl CoA pathway

Reductive hydroxypropionante pathway

Cathin-Bonion cycle

Figure 1 The distribution of autotrophic metabolic pathways among taxa within the three major domains of life(as inferred from 16S ribosomal RNA sequences (Pace, 1997)). Specific metabolic pathways are indicated.

Chemoautotrophy 187

in the ocean interior. Hence, this reaction is depen-dent on the chemical redox gradient between theventilating mantle plume and the ocean interiorthat thermodynamically favors oxidation of theplume gases. Maintaining such a gradient requiresa supply of energy, either externally, fromradiation (solar or otherwise), or internally, viaplanetary heat and tectonics, or both.

The overall contribution of chemoautotrophy inthe contemporary ocean to the formation oforganic matter is relatively small, accounting for,1% of the total annual primary production in thesea. However, this process is critical in couplingreduction of carbon to the oxidation of low-energysubstrates, and is essential for completion ofseveral biogeochemical cycles.

8.05.3 PHOTOAUTOTROPHY

The oxidation state of the ocean interior is aconsequence of a second energy source: light,which drives photosynthesis. Photosynthesis is aredox reaction of the general form:

2H2A þ CO2 þ light ! ðCH2OÞ þ H2O þ 2A

ð3Þ

where A is an atom, e.g., S. In this formulation,light is specified as a substrate, and a fraction ofthe light energy is stored as chemical bond energyin the organic matter. Organisms capable ofreducing inorganic carbon to organic matter byusing light energy to derive the source of reductantor energy are called photoautotrophs. Analyses ofgenes and metabolic sequences strongly suggestthat the machinery for capturing and utilizing lightas a source of energy to extract reductants wasbuilt on the foundation of chemoautotrophiccarbon fixation; i.e., the predecessors of photo-autotrophs were chemoautotrophs. The evolutionof a photosynthetic process in a chemoauto-troph forces consideration of both the selectiveforces responsible (why) and the mechanism ofevolution (how).

8.05.3.1 Selective Forces in the Evolutionof Photoautotrophy

Reductants for chemoautotrophs are generallydeep in the Earth’s crust. Vent fluids are producedin magma chambers connected to the Atheno-sphere. As such, the supply of vent fluids isvirtually unlimited. While the chemical disequili-bria between vent fluids and bulk seawaterprovides a sufficient thermodynamic gradient tocontinuously support chemoautotrophic meta-bolism in the contemporary ocean, in the earlyEarth the oceans would not have had a suffi-ciently large thermodynamic energy potential tosupport a pandemic outbreak of chemoautotrophy.

Moreover, magma chambers, vulcanism and ventfluid fluxes are tied to tectonic subduction andspreading regions, which are transient features ofEarth’s crust and hence only temporary habitatsfor chemoautotrophs. In the Archean and earlyProterozoic oceans, the chemoautotrophs wouldhave already been dispersed throughout the oceansby physical mixing and helping to colonize newvent regions. This same dispersion process wouldhave also helped ancestral chemoautotrophsexploit solar energy near the ocean surface.

Although the processes that selected the photo-synthetic reactions as the major energy transduc-tion pathway remain obscure, central hypotheseshave emerged based on our understanding of theevolution of Earth’s carbon cycle, the evolution ofphotosynthesis, biophysics, and molecular phylo-geny. Photoautotrophs are found in all three majorsuperkingdoms (Figure 1); however, there arevery few known Archea capable of this form ofmetabolism. Efficient photosynthesis requiresharvesting solar radiation, and hence the evolutionof a light harvesting system. While some Archeaand Bacteria use the pigment-protein rhodopsin,by far, the most efficient and ubiquitous lightharvesting systems are based on chlorins. Themetabolic pathway for the synthesis of porphyrinsand chlorins is one of the oldest in biologicalevolution, and is found in all chemoautotrophs(Xiong et al., 2000). Mulkidjanian and Junge(1997) proposed that the chlorin-based photosyn-thetic energy conversion apparatus originallyarose from the need to prevent UV radiationfrom damaging essential macromolecules such asnucleic acids and proteins. The UV excitationenergy could be transferred from the aromaticamino acid residues in the macromolecule to ablue absorption band of membrane-bound chlorinsto produce a second excited state which sub-sequently decays to the lower-energy excitedsinglet. This energy dissipation pathway can beharnessed to metabolism if the photochemicallyproduced, charge-separated, primary products areprevented from undergoing a back-reaction, butrather form a biochemically stable intermediatereductant. This metabolic strategy was selectedfor the photosynthetic reduction of CO2 tocarbohydrates, using reductants such as S22 orFe2þ, which have redox potentials that are toopositive to reduce CO2 directly.

The synthesis of reduced (i.e., organic) carbonand the oxidized form of the electron donorpermits a photoautotroph to use “respiratory”metabolism, but operate them in reverse. However,not all of the reduced carbon and oxidants remainaccessible to the photoautotrophs. In the oceans,cells tend to sink, carrying with them organiccarbon. The oxidation of Fe2þ forms insolubleFe3þ salts that precipitate. The sedimentation andsubsequent burial of organic carbon and Fe3þ


removes these components from the water column.Without replenishment, the essential reductantsfor anoxygenic photosynthesis would eventuallybecome depleted in the surface waters. Thus, thenecessity to regenerate reductants potentiallyprevented anoxygenic photoautotrophs from pro-viding the major source of fixed carbon on Earthfor eternity. Major net accumulation of reducedorganic carbon in Proterozoic sediments implieslocal depletion of reductants such as S22 and Fe2þ

from the euphotic zone of the ocean. Theselimitations almost certainly provided the evolu-tionary selection pressure for an alternativeelectron donor.

8.05.3.2 Selective Pressure in the Evolutionof Oxygenic Photosynthesis

H2O is a potentially useful biological reductantwith a vast supply on Earth relative to any redox-active solute dissolved in it. Liquid water contains,100 kmol of H atoms per m3, and, given.1018 m3 of water in the hydrosphere and cryo-sphere, .1020 kmol of reductant are potentiallyaccessible. Use of H2O as a reductant for CO2,however, requires a larger energy input than doesthe use of Fe2þ or S22. Indeed, to split water bylight energy requires 0.82 eV at pH 7 and298 K. Utilizing light at such high energy levelsrequired the evolution of a new photosyntheticpigment, chlorophyll a, which has a red (lowestsinglet) absorption band that is 200–300 nm blueshifted relative to bacteriochlorophylls. Moreover,stabilization of the primary electron acceptor toprevent a back-reaction necessitates thermody-namic inefficiency that ultimately requires twolight-driven reactions operating in series. This seq-uential action of two photochemical reactions isunique to oxygenic photoautotrophs and presum-ably involved horizontal gene transfer through oneor more symbiotic events (Blankenship, 1992).

In all oxygenic photoautotrophs, Equation (3)can be modified to:

2H2O þ CO2 þ light !Chl a

ðCH2OÞ þ H2O þ O2

ð4Þ

where Chl a is the pigment chlorophyll aexclusively utilized in the reaction. Equation (4)implies that somehow chlorophyll a catalyzes areaction or a series of reactions whereby lightenergy is used to oxidize water:

2H2O þ light !Chl a

4Hþ þ 4e þ O2 ð5Þ

yielding gaseous, molecular oxygen. Hiddenwithin Equation (5) are complex suites ofbiological innovations that have heretofore notbeen successfully mimicked in vitro by humans.At the core of the water splitting complex is a

quartet of manganese atoms, that sequentiallyextract electrons, one at a time, from 2H2O mole-cules, releasing gaseous O2 to the environment,and storing the reductants on biochemicalintermediates.

The photochemically produced reductants gene-rated by the reactions schematically outlined inEquation (5) are subsequently used in the fixation(fixation is an archaic term meaning to make non-volatile, as in the chemical conversion of a gas to asolid phase) of CO2 by a suite of enzymes that canoperate in vitro in darkness and, hence, theensemble of these reactions are called the darkreactions. At pH 7 and 25 8C, the formationof glucose from CO2 requires an investment of915 cal mol21. If water is the source of reductant,the overall efficiency for photosynthetic reductionof CO2 to glucose is ,30%; i.e., 30% of theabsorbed solar radiation is stored in the chemicalbonds of glucose molecules.

8.05.4 PRIMARY PRODUCTIVITYBY PHOTOAUTOTROPHS

When we subtract the costs of all othermetabolic processes by the chemoautotrophs andphotoautotrophs, the organic carbon that remainsis available for the growth and metabolic costs ofheterotrophs. This remaining carbon is called netprimary production (NPP) (Lindeman, 1942).From biogeochemical and ecological perspec-tives, NPP provides an upper bound for all othermetabolic demands in an ecosystem. If NPP isgreater than all respiratory consumption of theecosystem, the ecosystem is said to be netautotrophic. Conversely, if NPP is less than allrespiratory consumption, the system must eitherimport organic matter from outside its bounds, orit will slowly run down—it is net heterotrophic.

It should be noted that NPP and photosynthesisare not synonymous. On a planetary scale, theformer includes chemoautrophy, the latter doesnot. Moreover, photosynthesis per se does notinclude the integrated respiratory term for thephotoautotrophs themselves (Williams, 1993). Inreality, that term is extremely difficult to measuredirectly, hence NPP is generally approximatedfrom measurements of photosynthetic ratesintegrated over some appropriate length of time(a day, month, season, or a year) and respiratorycosts are either assumed or neglected.

8.05.4.1 What are Photoautotrophs?

In the oceans, oxygenic photoautotrophs area taxonomically diverse group of mostlysingle-celled, photosynthetic organisms that drift

Primary Productivity by Photoautotrophs 189

with currents. In the contemporary ocean, theseorganisms, called phytoplankton (derived fromGreek, meaning to wander), are comprised of,2 £ 104 species distributed among at leasteight taxonomic divisions or phyla (Table 2).

By comparison, higher plants are comprised of.2.5 £ 105 species, almost all of which arecontained within one class in one division. Thus,unlike terrestrial plants, phytoplankton are repre-sented by relatively few species but they are

Table 2 The taxonomic classification and species abundances of oxygenic photosynthetic organisms in aquatic andterrestrial ecosystems. Note that terrestrial ecosystems are dominated by relatively few taxa that are species rich,

while aquatic ecosystems contain many taxa but are relatively species poor.

Taxonomic group Known species Marine Freshwater

Empire: Bacteria (¼Prokaryota)Kingdom: EubacteriaSubdivision: Cyanobacteria (sensu strictu)

(¼Cyanophytes, blue-green algae)1,500 150 1,350

Subdivision: Chloroxybacteria(¼Prochlorophyceae)

3 2 1

Empire: EukaryotaKingdom: ProtozoaDivision: Euglenophyta 1,050 30 1,020Class: Euglenophyceae

Division: Dinophyta (Dinoflagellates)Class: Dinophyceae 2,000 1,800 200

Kingdom: PlantaeSubkingdom: Biliphyta

Division: GlaucocystophytaClass: Glaucocystophyceae 13

Division: RhodophytaClass: Rhodophyceae 6,000 5,880 120

Subkingdom: ViridiplantaeDivision: ChlorophytaClass: Chlorophyta 2,500 100 2,400

Prasinophyceae 120 100 20Ulvophyceae 1,100 1,000 100Charophyceae 12,500 100 12,400

Division: Bryophyta (mosses, liverworts) 22,000 1,000Division: Lycopsida 1,228 70Division: Filicopsida (ferns) 8,400 94Division: Magnoliophyta (flowering plants) (240,000)Subdivision: Monocotyledoneae 52,000 55 455Subdivision: Dicotyledoneae 188,000 391

Kingdom: ChromistaSubkingdom: ChlorechniaDivision: ChlorarachniophytaClass: Chlorarachniophyceae 3–4 3–4 0

Subkingdom: EuchromistaDivision: CrytophytaClass: Crytophyceae 200 100 100

Division: HaptophytaClass: Prymensiophyceae 500 100 400

Division: HeterokontaClass: Bacillariophyceae (diatoms) 10,000 5,000 5,000

Chrysophyceae 1,000 800 200Eustigmatophyceae 12 6 6Fucophyceae (brown algae) 1,500 1,497 3Raphidophyceae 27 10 17Synurophyceae 250 250Tribophyceae (Xanthophyceae) 600 50 500

Kingdom: FungiDivision: Ascomycontina (lichens) 13,000 15 20

Source: Falkowski (1997).


phylogenetically diverse. This deep taxonomicdiversity is reflected in their evolutionary historyand ecological function (Falkowski, 1997).

Within this diverse group of organisms, threebasic evolutionary lineages are discernable(Delwiche, 2000). The first contains all prokaryoticoxygenic phytoplankton, which belong to one classof bacteria, namely, the cyanobacteria. Cyanobac-teria are the only known oxygenic photoautotrophsthat existed prior to ,2.5 Gyr BP (Ga) (Lipps,1993; Summons et al., 1999). These prokaryotesnumerically dominate the photoautotrophic com-munity in contemporary marine ecosystems andtheir continued success bespeaks an extraordinaryadaptive capacity. At any moment in time, there are,1024 cyanobacterial cells in the contemporaryoceans. To put that into perspective, the number ofcyanobacterial cells in the oceans is two orders ofmagnitude more than all the stars in the sky.

The evolutionary history of cyanobacteria isobscure. The first microfossils assigned to thisgroup were identified in cherts from 3.1 Ga bySchopf (Schopf, 1993). Macroscopic stromabo-lites, which are generally of biological (oxygenicphotoautotrophic) origin, are first found in strata afew hundred million years younger. However,much of the fossil evidence provided by Schopf(e.g., Schopf, 1993) has been questioned (Brasieret al., 2002), and many researchers believe in a laterorigin. The origin of this group is critical toestablishing when net O2 production (and hence, anoxidized atmosphere) first occurred on the planet.Although photodissociation of H2O vapor couldhave provided a source of atmospheric O2 in theArchean, the UV absorption cross-section of O2

constrains the reaction, and theoretical calculationssupported by geochemical evidence suggest thatprior to ca. 2.4 Ga atmospheric O2 was less than1025 of the present level (Holland and Rye, 1998;Pavlov and Kasting, 2002). There was a lagbetween the first occurrence of oxygenic photo-synthesis and a global buildup of O2 possibly due tothe presence of alternative electron acceptors,especially Fe2þ and S22 in the ocean. Indeed, thedating of oxidation of Earth’s oceans and atmos-phere is, in large measure, based on analysis ofthe chemical precipitation of oxidized ironin sedimentary rocks (the “Great Rust Event(Holland and Rye, 1998) and the mass-independent(Farquhar et al., 2002) and mass-dependent(Habicht et al., 2002) fractionation of sulfurisotopes. The ensemble of these analyses indicatethat atmospheric oxygen rose sharply, fromvirtually insignificant levels, to between 1% and10% of the present atmospheric concentration overa 100 Myr period beginning ca. 2.4 Ga. Thus, theremay be as much as a 1 Gyr or as little as a 100 Myrgap between the origin of the first oxygenicphotoautotrophs and oxygenation of Earth’soceans and atmosphere.

All other oxygen-producing organisms in theocean are eukaryotic, i.e., they contain internalorganelles, including a nucleus, one or morechloroplasts, one or more mitochondria, and, insome cases, a membrane-bound storage compart-ment, the vacuole. Within the eukaryotes, we candistinguish two major groups, both of whichappear to have descended from a commonancestor thought to be the endosymbiotic appro-priation of a cyanobacterium into a heterotrophichost cell (Delwiche, 2000). The appropriatedcyanobacterium became a chloroplast.

8.05.4.1.1 The red and green lineages

In one group of eukaryotes, chlorophyll b wassynthesized as a secondary pigment; this groupforms the “green lineage,” from which all higherplants have descended. The green lineage playeda major role in oceanic food webs and the carboncycle from ca. 1.6 Ga until the end-Permianextinction, ,250 Ma (Lipps, 1993). Since thattime however, a second group of eukaryotes hasrisen to ecological prominence in the oceans; thatgroup is commonly called the “red lineage”(Figure 2). The red lineage is comprised of severalmajor phytoplankton divisions and classes, ofwhich the diatoms, dinoflagellates, haptophytes(including the coccolithophorids), and the chryso-phytes are the most important. All of these groupsare comparatively modern organisms; indeed, therise of dinoflagellates and coccolithophoridsapproximately parallels the rise of dinosaurs onland, while the rise of diatoms approximatelyparallels the rise of mammals in the Cenozoic.The burial and subsequent diagenesis of organiccarbon produced primarily by members of the redlineage in shallow seas in the Mesozoic eraprovide the carbon source for many of thepetroleum reservoirs that have been exploited forthe past century by humans.

8.05.4.2 Estimating Chlorophyll Biomass

As implied in Equation (4), given an abundanceof the two physical substrates, CO2 and H2O,primary production is, to first order, dependent onthe concentration of the catalyst Chl a and light.The distribution of Chl a in the upper ocean can bediscerned from satellite images of ocean color.The physical basis of the measurement isstraightforward; making the measurements istechnically challenging. Imagine two small par-cels of water that are adjacent to each other. Asphotons from the sun enter the water column, theyare either absorbed or scattered. Water itselfabsorbs red wavelengths of light, at shorterwavelengths of the visible spectrum, light is not


as efficiently absorbed. However, because watermolecules can randomly move from one adjacentparcel to another, there are continuous minorchanges in density and hence in the refractiveindex of the water parcels. These minor changes inrefractive index lead to incoherence in the

downwelling light stream. The incoherence, inturn, increases the probability of photon scattering(a process called “fluctuation density scattering”),such that light in the shorter wavelengths is morelikely to be scattered back to space (Einstein,1910; Morel, 1974). If the ocean contained sterile,

Figure 2 The basic pathway leading the evolution of eukaryotic algae. The primary symbiosis of a cyanobacteriumwith a apoplastidic host gave rise to both chlorophyte algae and red algae. The chlorophyte line, through secondarysymbioses, gave rise to the “green” line of algae, one division of which was the predecessor of all higher plants.Secondary symbioses in the red line with various host cells gave rise to all the chromophytes, including diatoms,

cryptophytes, and haptophytes (after Delwiche, 2000).


pure seawater, an observer looking at the surfacefrom space would see the oceans as blue.However, chlorophyll a has a prominent absorp-tion band in the blue portion of the spectrum.Hence, in the presence of chlorophyll some of thedownwelling and upwelling photons from the sunare absorbed by the phytoplankton themselvesand the ocean becomes optically darker. Aschlorophyll concentrations increase even further,blue wavelengths are largely eliminated fromthe outbound reflectance, and the ocean appearsoptically dark green (Morel, 1988).

8.05.4.2.1 Satellite based algorithms forocean color retrievals

Empirically, satellite sensors that measureocean color utilize a number of wavelengths. Inaddition to the blue and green region, red and far-red spectra are determined to derive correctionsfor scattering and absorption of the outbound orreflected radiation from the ocean by the atmos-phere. In fact, only a very small fraction (,5%) ofthe light leaving the ocean is observed by asatellite; the vast majority of the photons arescattered or absorbed in the atmosphere. How-ever, based on the ratio of blue-green light that isreflected from the ocean, estimates of photosyn-thetic pigments are derived. It should be pointedout that the blue-absorbing region of the spectrumis highly congested; it is virtually impossible toderive the fraction of absorption due solely tochlorophyll a as opposed to other photosyntheticpigments that absorb blue light. The estimation ofchlorophyll a is based on empirical regression ofthe concentration of the pigment to the total blue-absorbing pigments (Gordon and Morel, 1983).Water-leaving radiances ðLWÞ at specific wave-lengths are corrected for atmospheric scatteringand absorption, and the concentration of chloro-phyll is calculated from the ratios of blue andgreen light reflected from the water body. Thecalibration of the sensors is empirical andspecifically derived for individual satellites.Examples of such algorithms for five satellitesare given in Table 3.

One limitation of satellite images of oceanchlorophyll is that they do not provide informationabout the vertical distribution of phytoplankton.The water-leaving radiances visible to an observeroutside of the ocean are confined to the upper 20%of the euphotic zone (which is empirically definedas the depth to which 1% of the solar radiationpenetrates). In the open ocean there is almostalways a subsurface chlorophyll maximum that isnot visible to satellite ocean color sensors. Anumber of numerical models have been developedto estimate the vertical distribution of chlorophyllbased on satellite color data (Berthon and Morel,1992; Platt, 1986). The models rely on statisticalparametrizations and require numerous in situobservations to obtain typical profiles for a givenarea of the world ocean (Morel and Andre, 1991;Platt and Sathyendranath, 1988). In addition, largequantities of phytoplankton associated with thebottom of ice flows in both the Arctic andAntarctic are not visible to satellite sensors butdo contribute significantly to the primary pro-duction in the polar seas (Smith and Nelson,1990). Despite these deficiencies, the satellite dataallow high-resolution, synoptic observations ofthe temporal and spatial changes in phytoplanktonchlorophyll in relation to the physical circulationof the atmosphere and ocean on a global scale.

The global distribution of phytoplanktonchlorophyll in the upper ocean for winter andsummer, derived from a compilation of satelliteimages, is shown in Figure 3. To a first order, theimages reveal how the horizontal and temporaldistribution of phytoplankton is related to thephysical circulation of the oceans, especially themajor features of the basin-scale gyres. Forexample, throughout most of the central oceanbasins, between 308N and 308 S, phytoplanktonbiomass is extremely low, averaging 0.1–0.2 mgchlorophyll a m23 at the sea surface. In theseregions the vertical flux of nutrients is generallyextremely low, limited by eddy diffusion throughthe thermocline. Most of the chlorophyll biomassis associated with the thermocline. Because thereis no seasonal convective overturn in thislatitude band, there is no seasonal variation in

Table 3 Algorithms used to calculate chlorophyll a (C) from remote sensing reflectance. R is determined as themaximum of the values shown. Sensor algorithms are for Sea-viewing Wide Field of view Sensor (SeaWiFS), OceanColor and Temperature Scanner (OCTS), Moderate Resolution Imaging Spectroradiometer (MODIS), Coastal Zone

Color Scanner (CZCS), and Medium Resolution Imaging Spectrometer (MERIS).

Sensor Equation R

SeaWiFS/OC2 C¼10:0ð0:34123:001Rþ2:811R222:041R3Þ 2 0:04 490=555OCTS/OC4O C¼10:0ð0::40522:900Rþ1:690R220:530R321:144R4Þ 443.490.520=565MODIS/OC3M C¼10:0ð0:283022:753Rþ1:457R220:659R321:403R4Þ 443.490=550CZCS/OC3C C¼10:0ð0:36224:066Rþ5:125R222:645R320:597R4Þ 443.520=550MERIS/OC4E C¼10:0ð0:36822:814Rþ1:456R2þ0:768R321:292R4Þ 443.490.510=560SeaWiFS/OC4v4 C¼10:0ð0:36623:067Rþ1:930R2þ0:649R321:532R4Þ 443.490.510=555SeaWiFS/OC2v4 C¼10:0ð0:31922:336Rþ0:879R220:135R3Þ 2 0:071 490=555


phytoplankton chlorophyll. The chlorophyll con-centrations are slightly increased at the equator inthe Pacific and Atlantic Oceans, and south of theequator in the Indian Ocean. In the equatorialregions the thermocline shoals laterally as a resultof long-range wind stress at the surface (Pickardand Emery, 1990). The wind effectively piles upwater along its fetch, thereby inclining the uppermixed layer. This results in increased nutrientfluxes, shallower mixed layers, and higher chloro-phyll concentrations on the eastern end of theequatorial band, and decreased nutrient fluxes,deeper mixed layers, and lower chlorophyllconcentrations on the western end. This effectis most pronounced in the Pacific. The displace-ment of the band south of the equator in theIndian Ocean is primarily a consequence of basinscale topography.

At latitudes above ,308, a seasonal cycle inchlorophyll can occur (Figure 3). In the northernhemisphere, areas of high chlorophyll are foundin the open ocean of the North Atlantic in the springand summer. The southern extent and intensity ofthe North Atlantic phytoplankton bloom are notfound in the North Pacific. The North Atlantic

bloom is associated with deep vertical convectivemixing, which allows resupply of nutrients to theupper mixed layer of the ocean. This phenomenondoes not occur in the Pacific due to a strongervertical density gradient in that basin (driven by thehydrological cycle). The North Atlantic bloomleads to a flux of organic matter into the oceaninterior that is observed even at the seafloor.

In the southern hemisphere, phytoplanktonchlorophyll is generally lower at latitudessymmetrical with the northern hemisphere inthe corresponding austral seasons. For example,in the austral summer (January–March), phyto-plankton chlorophyll is slightly lower between308 S and the Antarctic ice sheets than in thenorthern hemisphere in July to September(Yoder et al., 1993).

8.05.4.3 Estimating Net Primary Production

8.05.4.3.1 Global models of net primaryproduction for the ocean

Using satellite data to estimate upper-ocean chlorophyll concentrations, satellite-based

Figure 3 Composite global images for winter and summer of upper-ocean chlorophyll concentrations (top panels)derived from satellite-based observations of ocean color, net primary production (middle panels) calculated based onthe algorithms of Behrenfeld and Falkowski (1997a), and export production (lower panels) calculated from the model

of Laws et al. (2000).


observations of incident solar radiation, atlases ofseasonally averaged sea-surface temperature, andmodels that incorporate a temperature responsefunction for photosynthesis, it is possible toestimate global net photosynthesis in the worldoceans (Antoine and Morel, 1996; Behrenfeld andFalkowski, 1997a; Longhurst et al., 1995).Although estimates vary between models, basedon how the parameters are derived, for illustrativepurposes we use a model based on empiricalparametrization of the daily integrated photosyn-thesis profiles as a function of depth. The physicaldepth at which 1% of irradiance incident on thesea surface remains is called the euphotic zone.This depth can be calculated from surfacechlorophyll concentrations, and defines the baseof the water column at which net photosynthesiscan be supported. Given such information, netprimary production can be calculated followingthe general equation:

PPeu ¼ Csat · Zeu · Pbopt · DL · F ð6Þ

where PPeu is daily net primary productionintegrated over the euphotic zone, Csat is thesatellite-based (upper water column; i.e., derivedfrom Table 3) chlorophyll concentration, Pb

opt isthe maximum daily photosynthetic rate within thewater column, Zeu is the depth of the euphoticzone, DL is the photoperiod (Behrenfeld andFalkowski, 1997a), and F is a function describingthe shape of the photosynthesis depth profile.

This general model can be both expanded(differentiated) and collapsed (integrated) withrespect to time and irradiance; however, the globalresults are fundamentally similar (Behrenfeld andFalkowski, 1997). The models predict that NPP inthe world oceans amounts to 40–50 Pg per annum(Figure 3 and Table 4).

In contrast to terrestrial ecosystems, the funda-mental limitation of primary production in theocean is not irradiance per se, but temperatureand the concentration of chlorophyll in the upperocean. The latter is a negative feedback; i.e., themore the chlorophyll in the water column,the shallower is the euphotic zone. Hence, todouble NPP requires nearly a fivefold increase inchlorophyll concentration.

8.05.4.4 Quantum Efficiency of NPP

The photosynthetically available radiation(400–700 nm) for the world oceans is 4.5 £ 1018

mol of photons per annum, which is .9.8 £ 1020

kJ yr21. The average energy stored by photosyn-thetic organisms amounts to ,39 kJ per gram ofcarbon fixed (Platt and Irwin, 1973). Given anannual net production of 40 Pg C for phytoplank-ton, and an estimated production of 4 Pg yr21 bybenthic photoautotrophs, the photosyntheticallystored radiation is equal to ,1.7 £ 1018 kJ yr21.The fraction of photosynthetically available solarenergy conserved by photosynthetic reactions in

Table 4 Annual and seasonal net primary production (NPP) of the major units of the biosphere.

Ocean NPP Land NPP

SeasonalApril–June 10.9 15.7July–September 13.0 18.0October–December 12.3 11.5January–March 11.3 11.2

BiogeographicOligotrophic 11.0 Tropical rainforests 17.8Mesotrophic 27.4 Broadleaf deciduous forests 1.5Eutrophic 9.1 Broadleaf and needleleaf forests 3.1Macrophytes 1.0 Needleleaf evergreen forests 3.1

Needleleaf deciduous forest 1.4Savannas 16.8Perennial grasslands 2.4Broadleaf shrubs with bare soil 1.0Tundra 0.8Desert 0.5Cultivation 8.0

Total 48.5 56.4

Source: Field et al. (1998). After Field et al. (1998). All values in GtC. Ocean color data are averages from 1978 to 1983. The land vegetationindex is from 1982 to 1990. Ocean NPP estimates are binned into three biogeographic categories on the basis of annual average Csat for each satellitepixel, such that oligotrophic ¼ Csat , 0:1 mg m23, mesotrophic ¼ 0:1 , Csat , 1 mg m23, and eutrophic ¼ Csat . 1 mg m23 (Antoine et al., 1996).This estimate includes a 1 GtC contribution from macroalgae (Smith, 1981). Differences in ocean NPP estimates between Behrenfeld andFalkowski (1997) and those in the global annual NPP for the biosphere and this table result from: (i) addition of Arctic and Antarctic monthly icemasks; (ii) correction of a rounding error in previous calculations of pixel area; and (iii) changes in the designation of the seasons to correspondwith Falkowski et al. (1998).


the world oceans amounts to 1.7 £ 1018 kJ/9.8 £ 1020 ¼ 0.0017 or 0.17%. Thus, on average,in the oceans, 0.0007 mol C is fixed per mole ofincident photons; this is equivalent to an effectivequantum requirement of 1,400 quanta per CO2

fixed. This value is less than 1% of the theoreticalmaximum quantum efficiency of photosynthesis;the relatively small realized efficiency is due to thefact that photons incident on the ocean surfacehave a small probability of being absorbed byphytoplankton before they are either absorbed bywater or other molecules (e.g., organic matter), orare scattered back to space.

8.05.4.4.1 Comparing efficiencies for oceanicand terrestrial primary production

The average chlorophyll concentration ofthe world ocean is 0.24 mg m23 and theaverage euphotic zone depth is 54 m; thus theaverage integrated chlorophyll concentration is,13 mg m22. Carbon to chlorophyll ratios ofphytoplankton typically range between 40:1 and100:1 by weight (Banse, 1977). Given the totalarea of the ocean of 3.1 £ 108 km2, the totalcarbon biomass in phytoplankton is 0.25–0.65 Pg.If NPP is ,40 Pg per annum, and assuming theocean is in steady state (a condition we willdiscuss in more detail), the living phytoplanktonbiomass turns over 60–150 times per year, whichis equivalent to a turnover time of 2–6 d. Incontrast, terrestrial plant biomass amounts to,600–1,000 Pg C, most of which is in the formof wood (Woodwell et al., 1978). Estimates ofterrestrial plant NPP are in the range of 50–65 PgC per annum, which gives an average turnovertime of ,12–20 yr (Field et al., 1998). Thus, theflux of carbon through aquatic photosyntheticorganisms is about 1,000-fold faster than terres-trial ecosystems, while the storage of carbon in thelatter is about 1,000-fold higher than the former.Moreover, the total photon flux to terrestrialenvironments amounts to ,2 £ 1018 mol yr21,which gives an effective quantum yield of,0.002. In other words, on average one CO2

molecule is fixed for every 500 incident photons.The results of these calculations suggest thatterrestrial vegetation is approximately three timesmore efficient in utilizing incident solar radiationto fix carbon than are aquatic photoautotrophs.This situation arises primarily from the relativepaucity of aquatic photoautotrophs in the oceanand the fact that they must compete with themedia (water) for light.

This comparison points out a fundamental diffe-rence between the two ecosystems in the context ofthe global carbon cycle. On timescales of decadesto centuries, carbon fixed in terrestrial ecosystemscan be temporarily stored in organic matter (e.g.,forests), whereas most of the carbon fixed by

marine phytoplankton is rapidly consumed bygrazers or sinks and is transferred from the surfaceocean to the ocean interior. Upon entering theocean interior, virtually all of the organic matteris oxidized by heterotrophic microbes, and in theprocess is converted back to inorganic carbon.Elucidating how this transfer occurs, what controlsit, how much carbon is transferred via this mecha-nism, on what timescales, and whether the processis in steady state was a major focus for research inthe latter portion of the twentieth century.

8.05.5 EXPORT, NEW AND “TRUE NEW”PRODUCTION

We can imagine that NPP produced by photo-autotrophs in the upper, sunlit regions of the ocean(the euphotic zone) is consumed in the samegeneral region by heterotrophs. In such a case, thebasic reaction given by Equation (4) is simplybalanced in the reverse direction due to respirationby heterotrophic organisms, and no organic matterleaves the ecosystem. This very simple “balancedstate” model, also referred to as the microbial loop(e.g., Azam, 1998), accounts for the fate of most ofthe organic matter in the oceans (and on time-scales of decades, terrestrial ecosystems as well).In marine ecology, this process is sometimescalled “regenerated production”; i.e., organicmatter produced by photoautotrophs is locallyregenerated to inorganic nutrients (CO2, NH4

þ,PO3

22) by heterotrophic respiration. It should benoted here that with the passage of organic matterfrom one level of a marine food chain to the next(e.g., from primary producer to heterotrophicconsumer), a metabolic “tax” must be paid in theform of respiration, such that the net metabolicpotential of the heterotrophic biomass is alwaysless than that of the primary producers. This doesnot mean, a priori, that photoautotrophic biomassis always greater, as heterotrophs may growslowly and accumulate biomass; however, asheterotrophs grow faster, their respiratory ratesmust invariably increase. The rate of production(i.e., the energy flux) of heterotrophic biomass isalways constrained by NPP.

Let us imagine a second scenario. Somefraction of the primary producers and/or hetero-trophs sink below a key physical gradient, such asa thermocline, and for whatever reason cannotascend back into the euphotic zone. If the watercolumn is very deep, sinking organic matter willmost likely be consumed by heterotrophicmicrobes in the ocean interior. The flux of organiccarbon from the euphotic zone is often calledexport production, a term coined by WolfgangBerger. Export production is an important conduitfor the exchange of carbon between the upperocean and the ocean interior (Berger et al., 1987).


This conduit depletes the upper ocean of inorganiccarbon and other nutrients essential for photosyn-thesis and the biosynthesis of organic matter. Inthe central ocean basins, export production is arelatively small fraction of total primary pro-duction, amounting to 5–10% of the total carbonfixed per annum (Dugdale and Wilkerson, 1992).At high latitudes and in nutrient-rich areas,however, diatoms and other large, heavy cellscan form massive blooms and sink rapidly. Insuch regions, export production can accountfor 50% of the total carbon fixation (Bienfang,1992; Campbell and Aarup, 1992; Sancetta et al.,1991; Walsh, 1983). The subsequent oxidationand remineralization of the exported produc-tion enriches the ocean interior with inorganiccarbon by ,200 mM in excess of that whichwould be supported solely by air–sea exchange(Figure 4 and Table 5). This enrichment is calledthe biological pump (Broecker et al., 1980;Sarmiento and Bender, 1994; Volk and Hoffert,1985). The biological pump is crucial to

maintaining the steady-state levels of atmosphericCO2 (Sarmiento et al., 1992; Siegenthaler andSarmiento, 1993).

8.05.5.1 Steady-state versus Transient State

The concepts of new, regenerated, and exportproduction are central to understanding manyaspects of the role of aquatic photosyntheticorganisms in biogeochemical cycles in the oceans.In steady state, the globally averaged fluxes ofnew nutrients must match the loss of the nutrientscontained in organic material. If this were not so,there would be a continuous depletion of nutrientsin the euphotic zone and photoautotrophic bio-mass and primary production would slowlydecline (Eppley, 1992). Thus, in the steady state,the sinking fluxes of organic nitrogen and theproduction of N2 by denitrifying bacteria mustequal the sum of the upward fluxes of inorganicnitrogen, nitrogen fixation, and the atmospheric

Figure 4 Vertical profiles of total dissolved inorganic carbon (TIC) in the ocean. Curve A corresponds to atheoretical profile that would have been obtained prior to the Industrial Revolution with an atmospheric CO2

concentration of 280 mmol mol21. The curve is derived from the solubility coefficients for CO2 in seawater, using atypical thermal and salinity profile from the central Pacific Ocean, and assumes that when surface water cools andsinks to become deep water it has equilibrated with atmospheric CO2. Curve B corresponds to the same calculatedsolubility profile of TIC, but in the year 1995, with an atmospheric CO2 concentration of 360 mmol mol21. Thedifference between these two curves is the integrated oceanic uptake of CO2 from anthropogenic emissions sincethe beginning of the Industrial Revolution, with the assumption that biological processes have been in steady state(and hence have not materially affected the net influx of CO2). Curve C is a representative profile of measured TICfrom the central Pacific Ocean. The difference between curve C and B is the contribution of biological processes tothe uptake of CO2 in the steady state (i.e. the contribution of the “biological pump” to the TIC pool.) (courtesy of

Doug Wallace and the World Ocean Circulation Experiment).

Export, New and “True New” Production 197

deposition of fixed nitrogen in the form of aero-sols (the latter is produced largely as a conse-quence of air pollution and, to a lesser extent, fromlightning).

8.05.6 NUTRIENT FLUXES

Primary producers are not simply sacks oforganic carbon. They are composed of six majorelements, namely, hydrogen, carbon, oxygen,nitrogen, phosphorus, and sulfur, and at least 54other trace elements and metals (Schlesinger,1997). In steady state, the export flux of organicmatter to the ocean interior must be coupled to theupward flux of several of these essential nutrients.The fluxes of nutrients are related to the elementalstoichiometry of the organic matter that sinksinto the ocean interior. This relationship, firstpointed out by Alfred Redfield in 1934, was basedon the chemistry of four of the major elementsin the ocean, namely, carbon, nitrogen, phos-phorus, and oxygen.

8.05.6.1 The Redfield Ratio

In the ocean interior, the ratio of fixed inorganicnitrogen (in the form of NO32) to PO4 in thedissolved phase is remarkably close to the ratios ofthe two elements in living plankton. Hence, itseemed reasonable to assume that the ratio of thetwo elements in the dissolved phase was the resultof the sinking and subsequent remineralization

(i.e., oxidation) of the elements in organic matterproduced in the open ocean. Further, as carbonand nitrogen in living organisms are largely foundin chemically reduced forms, while remineralizedforms are virtually all oxidized, the remineraliza-tion of organic matter was coupled to the depletionof oxygen. The relationship could be expressedstoichiometrically as

½106ðCH2OÞ16NH31PO224 � þ 138O2

! 106CO2 þ 122H2O þ 16NO23

þ PO224 þ 16Hþ� ð7Þ

Hidden within this balanced chemical formulationare biochemical redox reactions, which arecontained within specific groups of organisms.(In the oxidation of organic matter, there is someambiguity about the stoichiometry of O2/P.Assuming that the mean oxidation level of organiccarbon is that of carbohydrate (as is the case inEquation (9)), then the oxidation of that carbon isequimolar with O. Alternatively, some organicmatter may be more or less reduced thancarbohydrate, and therefore require more or lessO for oxidation. Note also that the oxidation ofNH3 to NO3

2 requires four atoms of O, and leads tothe formation of one H2O and one Hþ.) When thereactions primarily occur at depth, Equation (7) isdriven to the right, while when the reactionsprimarily occur in the euphotic zone, they aredriven to the left. Note that in addition to reducingCO2 to organic matter, formation of organicmatter by photoautotrophs requires reduction ofnitrate to the equivalent of ammonia. These twoforms of nitrogen are critically important inhelping to quantify “new” and export production.

8.05.7 NITRIFICATION

NO32 is produced via oxidation of NH3 by a

specific group of eubacteria, the nitrifiers, that areoblibate aerobes found primarily in the watercolumn. The oxidation of NH3 is coupled to thereduction of inorganic carbon to organic matter;hence nitrification is an example of a chemoauto-trophic process that couples the aerobic nitrogencycle to the carbon cycle. However, because thethermodynamic gradient is very small, the effi-ciency of carbon fixation by nitrifying bacteria islow and does not provide an ecologicallysignificant source of organic matter in the oceans.In the contemporary ocean, global CO2 fixation bymarine nitrifying bacteria only amounts to,0.2 Pg C per annum, or ,0.5% of marinephotoautotrophic carbon fixation.

There are two major sources of nutrients in theeuphotic zone. One is the local regeneration ofsimple forms of combined elements (e.g., NH4

þ,HPO4

22, SO422) resulting from the metabolic

Table 5 Export production and ef ratios calculatedfrom the model of Laws et al. (2000).

Export(GtC y21)

ef

Ocean basinPacific 4.3 0.19Atlantic 4.3 0.25Indian 1.5 0.15Antarctic 0.62 0.28Arctic 0.15 0.56Mediterranean 0.19 0.24Global 11.1 0.21

Total productionOligotrophic

(chl a , 0.1 mg m23)1.04 0.15

Mesotrophic(0.1 # chl a , 1.0 mg m23)

6.5 0.18

Eutrophic(chl a $ 1.0 mg m23)

3.6 0.36

Ocean depth0 2 100 m 2.2 0.31100 m–1 km 1.4 0.33.1 km 7.4 0.18


activity of metazoan and microbial degradation.The second is the influx of distantly produced,“new” nutrients, imported from the deep ocean,the atmosphere (i.e., nitrogen fixation, atmos-pheric pollution), or terrestrial runoff fromstreams, rivers, and estuaries (Dugdale andGoering, 1967). In the open ocean, these twosources can be usefully related to the form ofinorganic nitrogen assimilated by phytoplankton.Because biological nitrogen fixation is relativelylow in the ocean (see below) and nitrification inthe upper mixed layer is sluggish relative to theassimilation of nitrogen by photoautotrophs,nitrogen supplied from local regeneration isassimilated before it has a chance to becomeoxidized. Hence, regenerated nitrogen is primarilyin the form of ammonium or urea. In contrast, thefixed inorganic nitrogen in the deep ocean hassufficient time (hundreds of years) to becomeoxidized, and hence the major source of newnitrogen is in the form of nitrate. Using 15NH4 and15NO3 as tracers, it is possible to estimate thefraction of new nitrogen that fuels phytoplanktonproduction (Dugdale and Wilkerson, 1992). Thisapproach provides an estimate of both the up-ward flux of nitrate required to sustain the 15NO3

2

supported production, as well as the downwardflux of organic carbon, which is required tomaintain a steady-state balance (Dugdale andWilkerson, 1992; Eppley and Peterson, 1979).

8.05.7.1 Carbon Burial

On geological timescales, there is one import-ant fate for NPP, namely, burial in the sediments.By far the largest reservoir of organic matter onEarth is locked up in rocks (Table 1). Virtually allof this organic carbon is the result of the burial ofexported marine organic matter in coastal sedi-ments over literally billions of years of Earth’shistory. On geological timescales, the burial ofmarine NPP effectively removes carbon frombiological cycles, and places most (not all) of thatcarbon into the slow carbon cycle. A smallfraction of the organic matter escapes tectonicprocessing via the Wilson cycle and is perma-nently buried, mostly in continental rocks. Theburial of organic carbon is inferred not from directmeasurement, but rather from indirect means. Oneof the most common proxies used to derive burialon geological timescales is based on isotopicfractionation of carbonates. The rationale for thisanalysis is that the primary enzyme responsiblefor inorganic carbon fixation is ribulose 1,5-bisphosphate carboxylase/oxygenase (RuBisCO),which catalyzes the reaction between ribulose 1,5-bisphosphate and CO2 (not HCO3

2), to form twomolecules of 3-phosphoglycerate. The enzymestrongly discriminates against 13C, such that theresulting isotopic fractionation amounts to ,27‰

relative to the source carbon isotopic value. Theextent of the actual fractionation is somewhatvariable and is a function of carbon availabilityand of the transport processes for inorganic carboninto the cells, as well as the specific carboxylationpathway (Laws et al., 1997). However, regardlessof the quantitative aspects, the net effect of carbonfixation is an enrichment of the inorganic carbonpool in 13C, while the organic carbon produced isenriched in 12C.

8.05.7.2 Carbon Isotope Fractionation in OrganicMatter and Carbonates

The isotopic fractionation in carbonates mirrorsthe relative amount of organic carbon buried. It isgenerally assumed that the source carbon, fromvulcanism (the so-called “mantle” carbon) has anisotopic value of approximately 25‰. As massbalance must constrain the isotopic signatures ofcarbonate carbon and organic carbon with themantle carbon, then

forg ¼dw 2 dcarb

DB

ð8Þ

where forg is the fraction of organic carbon buried,dw is the average isotopic content of the carbonweathered, dcarb is the isotopic signature of thecarbonate carbon, and DB the isotopic differencebetween organic carbon and carbonate carbondeposited in the ocean. Equation (8) is a steady-state model that presumes the source of carbonfrom the mantel is constant over geological time.This basic model is the basis of nearly allestimates of organic carbon burial rates (Berneret al., 1983; Kump and Arthur, 1999).

Carbonate isotopic analyses reveal positiveexcursions (i.e., implying organic carbon burial)in the Proterozioc, and more modest excursionsthroughout the Phanaerozic (Figure 5). Burial oforganic carbon on geological timescales impliesthat export production must deviate from thesteady state on ecological timescales. Such adeviation requires changing one or more of(i) ocean nutrient inventories, (ii) the utilizationof unused nutrients in enriched areas, (iii) theaverage elemental composition of the organicmaterial, or (iv) the “rain” ratios of particulateorganic carbon to particulate inorganic carbon tothe seafloor.

8.05.7.3 Balance between Net PrimaryProduction and Losses

In the ecological theater of aquatic ecosystems,the observed photoautotrophic biomass at anymoment in time represents a balance between therate of growth and the rate of removal of thattrophic level. The burial of organic carbon in

Nitrification 199

Figure 5 Phanerozoic d13Ccarb record. The Jurassic through the Cenozoic record was generated from bulk sediment carbonates primarily from open ocean Atlantic Deep Sea DrillingProject boreholes; Lower Jurassic samples were used from the Mochras Borehole (Wales) (Katz et al., in review). Dashed intervals indicate data gaps. Singular spectrum analysis wasused to generate the Mesozoic–Cenozoic curve. The Paleozoic record was generated from brachiopods (Veizer et al., 1999; note that the timescale has been adjusted from the originalreference by interpolating between period boundaries). Data were averaged for each time slice to obtain the Paleozoic curve. The timescales of Berggren et al. (1995; Cenozoic),

Gradstein et al. (1995; Mesozoic), and GSA (Paleozoic) were used.

the lithosphere requires that the ecologicalbalance between NPP and respiration diverge;i.e., the global ocean must be net autotropic. Forsimplicity, we can express the time-dependentchange in photoautrophic biomass by a lineardifferential equation:

dP=dt ¼ ½P�ðm2 mÞ ð9Þ

where ½P� is photoautrophic biomass (e.g., organiccarbon), m is the specific growth rate (units of1/time), and m is the specific mortality rate (unitsof 1/time). In this equation, we have lumped allmortality terms, such as grazing and sinkingtogether into one term, although each of theseloss processes can be given explicitly (Banse,1994). Two things should be noted regardingEquation (9). First, m and m are independentvariables; i.e., changes in P can be independentlyascribed to one or the other process. Second, bydefinition, a steady state exists when dP=dt is zero.

8.05.7.4 Carbon Burial in the ContemporaryOcean

The burial of organic carbon in the modernoceans is primarily confined to a few regionswhere the supply of sediments from terrestrialsources is extremely high. Such regions includethe Amazon outfall and Indonesian mud belts. Incontrast, the oxidation of organic matter in theinterior of the contemporary ocean is extremelyefficient; virtually no carbon is buried in the deepsea. Similarly, on most continental margins,organic carbon that reaches the sediments isconsumed by microbes within the sediments,such that very little is actually buried in thecontemporary ocean (Aller, 1998). The solution toEquation (9) must be close to zero; and conse-quently, in the absence of human activities, theoxygen content of Earth’s atmosphere is veryclose to steady state and has been zso for tens, ifnot hundreds, of millions of years.

8.05.7.5 Carbon Burial in the PrecambrianOcean

In contrast, carbon burial in the Proterozoicocean must have occurred as oxygen increased inthe atmosphere; i.e., the global solution toEquation (9) must have been .0. Photoauto-trophic biomass could have increased until someelement became limiting. Thus, the originalfeedback between the production of photoauto-trophic biomass in the oceans and the atmosphericcontent of oxygen was determined by an elementthat limited the crop size of the photoautotrophs inthe Archean or Proterozoic ocean. What was thatelement, and why did it become limiting?

8.05.8 LIMITING MACRONUTRIENTS

A general feature of aquatic environments isthat because the oxidation of organic nutrients totheir inorganic forms occurs below the euphoticzone where the competing processes of assimila-tion of nutrients by photoautotrophs do not occur,the pools of inorganic nutrients are much higher atdepth. As the only natural source of photosynthe-tically active radiation is the sun, the gradients oflight and nutrients are from opposite directions.Thermal or salinity differences in the surfacelayers produce vertical gradients in density thateffectively retard the vertical fluxes of solublenutrients from depth. Thus, in the surface layers ofa stratified water column, nutrients becomedepleted as the photoautotrophs consume themat rates exceeding their rate of vertical supply.Indeed, throughout most of the world oceans, theconcentrations of dissolved inorganic nutrients,especially fixed inorganic nitrogen and phosphate,are exceedingly low, often only a few nM. One orthe other of these nutrients can limit primaryproduction. However, the concept of limitationrequires some discussion.

8.05.8.1 The Two Concepts of Limitation

The original notion of limitation in ecology wasrelated to the yield of a crop. A limiting factor wasthe substrate least available relative to therequirement for synthesis of the crop (Liebig,1840). This concept formed a strong underpinningof agricultural chemistry and was used to designthe elemental composition of fertilizers forcommercial crops. This concept subsequentlywas embraced by ecologists and geochemists asa general “law” (Odum, 1971).

Nutrients can also limit the rate of growth ofphotoautotrophs (Blackman, 1905; Dugdale,1967). Recall that if organisms are in balancedgrowth, the rate of uptake of an inorganic nutrientrelative to the cellular concentration of the nutrientdefines the growth rate (Herbert et al., 1956). Theuptake of inorganic nutrients is a hyperbolicfunction of the nutrient concentration and can beconveniently described by a hyperbolic expressionof the general form

V ¼ ðVmaxðU;…ÞÞ=ðKs þ ðU;…ÞÞ ð10Þ

where V is the instantaneous rate of nutrientuptake, Vmax is the maximum uptake rate, (U;…)represent the substrate concentration of nutrient U,etc., and Ks is the concentration supporting halfthe maximum rate of uptake (Dugdale, 1967;Monod, 1942). There can be considerablevariation between species with regard to Ks andVmax values and these variations are potential

Limiting Macronutrients 201

sources of competitive selection (Eppley et al.,1969; Tilman, 1982).

It should be noted that Liebig’s notion oflimitation was not related a priori to the intrinsicrate of photosynthesis or growth. For example,photosynthetic rates can be (and often are) limitedby light or temperature. The two concepts oflimitation (yield and rate) are often not understoodcorrectly: the former is more relevant to biogeo-chemical cycles, the latter is more critical toselection of species in ecosystems.

8.05.9 THE EVOLUTION OF THE NITROGENCYCLE

Globally, nitrogen and phosphorus are the twoelements that immediately limit, in a Liebig sense,the biologically mediated carbon assimilation inthe oceans by photoautotrophs. It is frequentlyargued that since N2 is abundant in both the oceanand the atmosphere, and, in principle, can bebiologically reduced to the equivalent of NH3 byN2-fixing cyanobacteria, nitrogen cannot be limit-ing on geological timescales (Barber, 1992;Broecker et al., 1980; Redfield, 1958). Therefore,phosphorus, which is supplied to the ocean by theweathering of continental rocks, must ultimatelylimit biological productivity. The underlyingassumptions of these tenets should, however, beconsidered within the context of the evolution ofbiogeochemical cycles.

By far, the major source of fixed inorganicnitrogen for the oceans is via biological nitrogenfixation. Although in the Archean atmosphere,electrical discharge or bolide impacts may havepromoted NO formation from the reactionbetween N2 and CO2, the yield for these reactionsis extremely low. Moreover, atmospheric NH3

would have photodissociated from UV radiation(Kasting, 1990), while N2 would have been stable(Kasting, 1990; Warneck, 1988). Biological N2

fixation is a strictly anaerobic process (Postgate,1971), and the sequence of the genes encoding thecatalytic subunits for nitrogenase is highly con-served in cyanobacteria and other eubacteria,strongly suggesting a common ancestral origin(Zehr et al., 1995). The antiquity and homology ofnitrogen fixation capacity also imply that fixedinorganic nitrogen was scarce prior to theevolution of diazotrophic organisms; i.e., therewas strong evolutionary selection for nitrogenfixation in the Archean or early Proterozoicperiods. In the contemporary ocean, N2 is stillcatalyzed solely by prokaryotes, primarily cyano-bacteria (Capone and Carpenter, 1982).

While apatite and other calcium-based andsubstituted solid phases of phosphate mineralsprecipitated in the primary formation of crustalsediments, secondary reactions of phosphate with

aluminum and transition metals such as iron aremediated at either low salinity, low pH, or highoxidation states of the cations (Stumm andMorgan, 1981). Although these reactions wouldreduce the overall soluble phosphate concen-tration, the initial condition of the Archeanocean probably had a fixed low N:P ratio in thedissolved inorganic phase. As N2 fixation pro-ceeded, that ratio would have increased with abuildup of ammonium in the ocean interior. Theaccumulation of fixed nitrogen in the oceanswould continue until the N:P ratio of the inorganicelements reached equilibrium with the N:P ratio ofthe sedimenting particulate organic matter (POM).Presumably, the latter ratio would approximatethat of extant, nitrogen-fixing marine cyanobac-teria, which is ,16:1 by atoms (Copin-Montegutand Copin-Montegut, 1983; Redfield, 1958;Quigg et al., 2003) or greater (Letelier et al.,1996) and would ultimately be constrained bythe availability of phosphate (Falkowski, 1997;Tyrell, 1999).

The formation of nitrate from ammonium bynitrifying bacteria requires molecular oxygen;hence, nitrification must have evolved followingthe formation of free molecular oxygen in theoceans by oxygenic photoautotrophs. Therefore,from a geological perspective, the conversion ofammonium to nitrate probably proceeded rapidlyand provided a substrate, NO3

2, that eventuallycould serve both as a source of nitrogen forphotoautotrophs and as an electron acceptor for adiverse group of heterotrophic, anaerobic bacteria,the denitrifiers.

In the sequence of the three major biologicalprocesses that constitute the nitrogen cycle,denitrification must have been the last to emerge.This process, which permits the reduction of NO3

2

to (ultimately) N2, occurs in the modern ocean inthree major regions, namely, continental marginsediments, areas of restricted circulation such asfjords, and oxygen minima zones of perenniallystratified seas (Christensen et al., 1987; Codispotiand Christensen, 1985; Devol, 1991; Nixon et al.,1996). In all cases, the process requires hypoxic oranoxic environments and is sustained by highsinking fluxes of organic matter. Denitrificationappears to have evolved independently severaltimes; the organisms and enzymes responsible forthe pathway are highly diverse from a phylogeneticand evolutionary standpoint.

With the emergence of denitrification, the ratioof fixed inorganic nitrogen to dissolved inorganicphosphate in the ocean interior could only bedepleted in nitrogen relative to the sinking flux ofthe two elements in POM. Indeed, in all of themajor basins in the contemporary ocean, the N:Pratio of the dissolved inorganic nutrients in theocean interior is conservatively estimated at 14.7


by atoms (Fanning, 1992) or less (Anderson andSarmiento, 1994) (Figure 6).

There are three major conclusions that may bedrawn from the foregoing discussion:

(i) Because the ratio of the sinking flux ofparticulate organic nitrogen and particulate phos-phorus exceeds the N:P ratio of the dissolved poolof inorganic nutrients in the ocean interior, theaverage upward flux of inorganic nutrients must beslightly enriched in phosphorus relative to nitrogenas well as to the elemental requirements of thephotoautotrophs (Gruber and Sarmiento, 1997;Redfield, 1958). Hence, although there are someexceptions (Kromer, 1995; Wu et al., 2000),dissolved, inorganic fixed nitrogen generally limitsprimary production throughout most of the world’soceans (Barber, 1992; Falkowski et al., 1998).

(ii) The N:P ratio of the dissolved pool ofinorganic nutrients in the ocean interior wasestablished by biological processes, not viceversa (Redfield et al., 1963, 1934). The elementalcomposition of marine photoautotrophs has beenconserved since the evolution of the eukaryotic

phytoplankton (Lipps, 1993). The Redfield N:Pratio of 16:1 for POM (Codispoti, 1995; Copin-Montegut and Copin-Montegut, 1983; McElroy,1983; Redfield et al., 1963, 1958) is an upperbound, which is not observed for the two elementsin the dissolved inorganic phase in the oceaninterior. The deficit in dissolved inorganic fixednitrogen relative to soluble phosphate in the oceanrepresents a slight imbalance between nitrogenfixation and denitrification on timescales of,103–104 yr (Codispoti, 1995).

(iii) If dissolved inorganic nitrogen rather thanphosphate limits productivity in the oceans, then itfollows that the ratio of nitrogen fixation/denitri-fication plays a critical role in determining the netbiologically mediated exchange of CO2 betweenthe atmosphere and ocean (Codispoti, 1995).

8.05.10 FUNCTIONAL GROUPS

As we have implied throughout the foregoingdiscussion, the biologically mediated fluxes of

Figure 6 (a)–(c) Vertical profiles of NO32/HPO4

32 ratios in each of the three major ocean basins. The data weretaken from the GEOSECS database. In all three basins, the N:P ratio converges on an average value that issignificantly lower than the 16:1 ratio predicted by Redfield. The deficit in nitrogen relative to phosphorus ispresumed to be a result of denitrification. Note that in the upper 500 m of the water column, NO3

2/HPO432 ratios

generally decline except in a portion of the Atlantic that corresponds to the eastern Mediterranean.

Functional Groups 203

elements between the upper ocean and the oceaninterior are critically dependent upon key groupsof organisms. Fluxes between the atmosphere andocean, as well as between the ocean and thelithosphere, are mediated by organisms thatcatalyze phase state transitions from either gas tosolute/solid or from solute to solid/gas phases. Forexample, autotrophic carbon fixation convertsgaseous CO2 to a wide variety of organic carbonmolecules, virtually all of which are solid ordissolved solids at physiological temperatures.Respiration accomplishes the reverse. Nitrogenfixation converts gaseous N2 to ammonium andthence to organic molecules, while denitrificationaccomplishes the reverse. Calcification convertsdissolved inorganic carbon and calcium to solidphase calcite and aragonite, whereas silicificationconverts soluble silicic acid to solid hydratedamorphous opal. Each of these biologicallycatalyzed processes is dependent upon specificmetabolic sequences (i.e., gene families encodinga suite of enzymes) that evolved over hundreds ofmillions of years of Earth’s history, and have, overcorresponding periods, led to the massive accumu-lation of oxygen in the atmosphere, and opal,carbonates, and organic matter in the lithosphere.Presumably, because of parallel evolution as wellas lateral gene transfer, these metabolic sequenceshave frequently co-evolved in several groups oforganisms that, more often than not, are notclosely related from a phylogenetic standpoint(Falkowski, 1997). Based on their biogeochemicalmetabolism, these homologous sets of organismsare called functional groups or biogeochemicalguilds; i.e., organisms that are related throughcommon biogeochemical processes rather than acommon evolutionary ancestor affiliation.

8.05.10.1 Siliceous Organisms

In the contemporary ocean, the export ofparticulate organic carbon from the euphoticzone is highly correlated with the flux ofparticulate silicate. Most of the silicate flux is aconsequence of precipitation of dissolved ortho-silicic acid by diatoms to form amorphous opalthat makes up the cell walls of these organisms.These hard-shelled cell walls presumably help theorganisms avoid predation, or if ingested, increasethe likelihood of intact gut passage through somemetazoans (Smetacek, 1999). In precipitatingsilicate, diatoms simultaneously fix carbon.Upon depleting the euphotic zone of nutrients,the organisms frequently sink en masse, and whilesome are grazed en route, many sink as intactcells. Ultimately, either fate leads to the gravi-tationally driven export flux of particulate organiccarbon into the ocean interior.

Silica is supplied to the oceans from the weath-ering of continental rocks. Because of precipitation

by silicious organisms, however, the ocean isrelatively depleted in dissolved silica. Althoughdiatom frustules (their silicified cell walls) tend todissolve and are relatively poorly preserved inmarine sediments, enough silica is buried to keepthe seawater undersaturated, throughout the ocean.As the residence time of silica in the oceans is,104 yr (i.e., about an order of magnitude longerthan the mean deep-water circulation), one can getan appreciation for the silicate demands andregeneration rates by following the concentrationgradients of dissolved silica along isopycnals.While these demands are generally attributed todiatoms, radiolarians (a group of nonphotosyn-thetic, heterotrophic protists with silicious teststhat are totally unrelated to diatoms) are notuncommon, and radiolarian shells are abundant inthe sediments of Southern Ocean. Silica is alsoprecipitated by various sponges and other protists.As a functional group, the silicate precipitators areidentified by their geochemical signatures in thesediments and in the silica chemistry of the oceans.

8.05.10.2 Calcium Carbonate Precipitation

Like silica precipitation, calcium carbonate isnot confined to a specific phylogenetically distinctgroup of organisms, but evolved (apparentlyindependently) several times in marine organisms.Carbonate sediments blanket much of the Atlanticbasin, and are formed from the shells of bothcoccolithophorids and foraminifera (Milliman,1993). (In the Pacific, the carbon compensationdepth is generally higher than the bottom, andhence, in that basin carbonates tend to dissolverather than become buried.) As the crystalstructure of the carbonates in both groups iscalcite (as opposed to the more diageneticallysusceptible aragonite), the preservation of theseminerals and their co-precipitating trace elementsprovides an invaluable record of ocean history.Although on geological timescales huge amountsof carbon are removed from the atmosphere andocean and stored in the lithosphere as carbonates,on ecological timescales, carbonate formationleads to the formation of CO2. This reaction canbe summarized by the following:

2HCO23 þ Ca2þ ! CaCO3 þ CO2 þ H2O

ð11Þ

Unlike silicate precipitation, calcium carbonateprecipitation leads to strong optical signatures thatcan be detected both in situ and remotely (Balchet al., 1991; Holligan and Balch, 1991). The basicprinciple of detection is the large, broadband (i.e.,“white”) scattering cross-sections of calcite. Thehigh scattering cross-sections are detected bysatellites observing the upper ocean as relativelyhighly reflective properties (i.e., a “bright” ocean).


Using this detection scheme, one can reconstructglobal maps of planktonic calcium carbonateprecipitating organisms in the upper ocean.In situ analysis can be accompanied by opticalrotation properties (polarization) to discriminatecalcite from other scattering particles. In situ pro-files of calcite can be used to construct the verticaldistribution of calcium carbonate-precipitatingplanktonic organisms that would otherwise notbe detected by satellite remote sensing becausethey are too deep in the water column.

Over geological time, the relative abundancesof key functional groups change. For example,relative coccolithophorid abundances generallyincreased through the Mesozoic, and underwent aculling at the Cretaceous Tertiary (K/T) boundary,followed by a general waning throughout theCenozoic. The changes in the coccolithophoridabundances appear to follow eustatic sea-levelvariations, suggesting that transgressions lead tohigher calcium carbonate deposition. In contrast,diatom sedimentation increases with regressionsand, since the K/T impact, diatoms have generallyreplaced coccolithophorids as ecologically import-ant eukaryotic phytoplankton. On much finertimescales during the Pleistocene, it would appearthat interglacial periods favor coccolithophoridabundance, while glacial periods favor diatoms.The factors that lead to glacial– interglacialvariations between these two functional groupsare relevant to elucidating their distributions in thecontemporary ecological setting of the ocean(Tozzi, 2001).

8.05.10.3 Vacuoles

In addition to a silicic acid requirement,diatoms, in contrast to dinoflagellates and cocco-lithophores, have evolved a nutrient storagevacuole (Raven, 1987). The vacuole, whichoccupies ,35% of the volume of the cell, canretain high concentrations of nitrate and phos-phate. Importantly, ammonium cannot be (or isnot) stored in a vacuole. The vacuole allowsdiatoms to access and hoard pulses of inorganicnutrients, thereby depriving potentially competinggroups of these essential resources. Consequently,diatoms thrive best under eutrophic conditions andin turbulent regions where nutrients are suppliedwith high pulse frequencies.

The competition between diatoms and cocco-lithophorids can be easily modeled by a resourceacquisition model based on nutrient uptake(Equation (9)). In such a model, diatoms domi-nate under highly turbulent conditions, whentheir nutrient storage capacity is maximallyadvantageous, while coccolithophorids dominateunder relatively quiescent conditions (Tozziet al., 2003).

The geological record during the Pleistocenereveals a periodicity of opal/calcite depositioncorresponding to glacial/interglacial periods. Suchalterations in mineral deposition are probablyrelated to upper ocean turbulence; i.e., thesedimentary record is a “fax” machine of mixing(Falkowski, 2002). Glacial periods appear to becharacterized by higher wind speeds and astronger thermal contrast between the equatorand the poles. These two factors would, inaccordance with the simple nutrient uptakemodel, favor diatoms over coccolithophores.During interglacials, more intense ocean stratifi-cation, weaker winds, and a smaller thermalcontrast between the equator and the poleswould tend to reduce upper-ocean mixing andfavor coccolithophores (Iglesias-Rodriguez et al.,2002). While other factors such as silica avail-ability undoubtedly also influenced the relativesuccess of diatoms and coccolithophores on thesetimescales, we suggest that the climatically forcedcycle, played out on timescales of 40 kyr and100 kyr (over the past 1.9 Myr), can be understoodas a long-term competition that never reachesan exclusion equilibrium condition (Falkowskiet al., 1998).

Can the turbulence argument be extended toeven longer timescales to account for the switch inthe dominance from coccolithophorids to diatomsin the Cenozoic? The fossil record of diatoms inthe Mesozoic is obscured by problems of preser-vation; however several species are preserved inthe late Jurassic (Harwood and Nikolaev, 1995),suggesting that the origins were in the earlyJurassic or perhaps as early as the Triassic. It isclear, however, that this group did not contributenearly as much to export production duringMesozoic times. We suggest that the ongoingsuccessional displacement of coccolithophores bydiatoms in the Cenozoic is, to first order, driven bytectonics (i.e., the Wilson cycle). The Mesozoicperiod was relatively warm and was characterizedby a two-cell Hadley circulation, with obliquitygreater than 378, resulting in a thoroughly mixedatmosphere with nearly uniform temperaturesover the surface of the Earth. The atmosphericmeridional heat transport decreased the latitudinalthermogradients; global winds and ocean circula-tion were both sluggish (Huber et al., 1995). Thisrelatively quiescent period of Earth’s history wasideal for coccolithophorids. Following the K/Timpact, and more critically, the onset of polar icecaps about 32 Ma, the Hadley circulation changeddramatically. Presently, there are six Hadley cells,and the atmosphere has become drier. The netresult is more intense thermohaline circulation,greater wind mixing and decreased stability(Barron et al., 1995; Chandler et al., 1992).Associated with this decreased stability is the riseof the diatoms.

Functional Groups 205

Over the past 50 Myr, both carbon and oxygenisotopic records in fossil foraminifera suggest thatthere has been a long-term depletion of CO2 in theocean–atmosphere system and a decrease intemperature in the ocean interior. The result hasbeen increased stratification of the ocean, whichhas, in turn, led to an increased importance forwind-driven upwelling and mesoscale eddy tur-bulence in providing nutrients to the euphoticzone. The ecological dominance of diatoms undersporadic mixing conditions suggests that theirlong-term success in the Cenozoic reflects anincrease in event-scale turbulent energy dissipa-tion in the upper ocean. But, was the Wilson cyclethe only driver?

Although weathering of siliceous minerals byCO2 (the so-called “Urey reactions”; but see(Berner and Maasch, 1996)) contributed to thelong-term flux of silica to the oceans (Berner,1990), and potentially fostered the radiation ofdiatoms in the Cenozoic, by itself, orogeny cannotexplain the relatively sharp increase in diatoms atthe Eocene/Oligocene boundary. Indeed, theseawater strontium isotope record does notcorrespond with these radiations in diatoms(Raymo and Ruddiman, 1992). We must lookfor other contributing processes.

Shortly after, or perhaps coincident withPaleocene thermal maximum (55 Ma), was a risein true grasses (Retallack, 2001). This group,which rapidly radiated in the Eocene, rose toprominence in Oligocene, a period coincidentwith a global climatic drying. During this period,however, there was a rapid co-evolution ofgrazing ungulates that displaced browsers (Janisand Damuth, 1990). Grasses contain up to 10%dry weight of silica, which forms micromineraldeposits in the cell walls; phytoliths (Conley,2002). Indeed, the selection of hypsodont (highcrown) dentition in ungulates from the brachydont(leaf eating) early-appearing browsing mammals,coincides with the widespread distribution ofphytoliths and grit in grassland forage. It istempting to suggest that the rise of grazingungulates, which spurred the radiation of grasses,was, in effect, a biologically catalyzed silicateweathering process. The deep-root structure ofEocene grasses certainly facilitated silicate mobil-ization into rivers and groundwaters (Conley,2002). Additionally, upon their annual death anddecay, the phytoliths of many temperate grassesare potentially transported to the oceans via wind.

The feedback between the co-evolution ofmammals and grasses and the supply of silicatesto the ocean potentially explains the rapidradiation of diatoms, and their continued dom-inance in the Cenozoic. There is another potentialfeedback at play, however, which “locked in”the diatom preeminence. It is likely that theincrease in diatom dominance, and the associated

increase in the efficiency of carbon burial, playeda key role in decreasing atmospheric CO2 over thepast 32 Myr. That biological selection mayinfluence climate is clearly controversial; howe-ver, the trends in succession between taxa ontimescales of tens of millions of years, and cyclesin dominance on shorter geological timescales begfor explanation.

8.05.11 HIGH-NUTRIENT,LOW-CHLOROPHYLLREGIONS—IRON LIMITATION

On ecological timescales, the biologicallymediated net exchange of CO2 between theocean and atmosphere is limited by nutrientsupply and the efficiency of nutrient utilizationin the euphotic zone. There are three major areasof the world ocean where inorganic nitrogen andphosphate are in excess throughout the year, yetthe mixed-layer depth appears to be shallowerthan the critical depth; these are the easternequatorial Pacific, the subarctic Pacific, andSouthern (i.e., Antarctic) Oceans. In the subarcticNorth Pacific, it has been suggested that there is atight coupling between phytoplankton productionand consumption by zooplankton (Miller et al.,1991). This grazer-limited hypothesis has beenused to explain why the phytoplankton in theNorth Pacific do not form massive blooms in thespring and summer like their counterparts inthe North Atlantic (Banse, 1992). In the mid-1980s, however, it became increasingly clear thatthe concentration of trace metals, especially iron,was extremely low in all three of these regions(Martin, 1991). Indeed, in the eastern equatorialPacific, for example, the concentration of solubleiron in the euphotic zone is only 100–200 pM.Although iron is the most abundant transitionmetal in the Earth’s crust, in its most commonlyoccurring form, Fe3þ, it is virtually insoluble inseawater. The major source of iron to the euphoticzone is Aeolian dust, originating from continentaldeserts. In the three major areas of the worldoceans with high inorganic nitrogen in the surfacewaters and low chlorophyll concentrations, theflux of Aeolian iron is extremely low (Duce andTindale, 1991). In experiments in which iron wasartificially added on a relatively large scale to thewaters in the equatorial Pacific, Southern Ocean,and subarctic Pacific, there were rapid anddramatic increases in photosynthetic energy con-version efficiency and phytoplankton chlorophyll(Abraham et al., 2000; Behrenfeld et al., 1996;Kolber et al., 1994; Tsuda et al., 2003). Beyonddoubt, NPP and export production in all threeregions are limited by the availability of a singlemicronutrient—iron.


8.05.12 GLACIAL–INTERGLACIALCHANGES IN THE BIOLOGICALCO2 PUMP

In the modern (i.e., interglacial) ocean, twomajor factors affect iron fluxes. First, changes inland-use patterns and climate over the past severalthousand years have had, and continue to have,marked effects on the areal distribution and extentof deserts. At the height of the Roman Empiresome 2,000 years ago, vast areas of North Africawere forested, whereas today these same areas aredesert. These changes were climatologicallyinduced. Similarly, the Gobi Desert in NorthCentral Asia has increased markedly in moderntimes. The flux of Aeolian iron from the SaharaDesert fuels photosynthesis for most of the NorthAtlantic Ocean; that from the Gobi is depositedover much of the North Pacific (Duce and Tindale,1991). The primary source of iron for the SouthernOcean is Australia, but the prevailing wind vectorsconstrain the delivery of the terrestrial dust tothe Indian Ocean. Consequently, the SouthernOcean is iron limited in the modern epoch(Martin, 1990).

The second factor in this climatological feed-back is that the major wind vectors are driven byatmosphere–ocean heat gradients. Changes inthermal gradients between the equator and poleslead to changes in wind speed and direction. Windvectors prior to glaciations appear to havesupported high fluxes of iron to the SouthernOcean, thereby presumably stimulating phyto-plankton production, the export of carbon todepth, and the drawdown of atmospheric CO2

appears to have accompanied glaciations in therecent geological past (Berger, 1988).

8.05.13 IRON STIMULATION OF NUTRIENTUTILIZATION

The enhancement of net export production(i.e., “true” new production) requires the additionof a limiting nutrient to the ocean, an increase inthe efficiency of utilization of preformed nutrientsin the upper ocean, and/or a change in theelemental stoichiometry of primary producers(Falkowski et al., 1998; Sarmiento and Bender,1994). Indeed, an analysis of ice cores fromAntarctica, reconstruction of Aeolian iron deposi-tions and concurrent atmospheric CO2 concen-trations over the past 4.2 £ 105 yr (spanning fourglacial–interglacial cycles) suggests that, wheniron fluxes were high, CO2 levels were low andvice versa (Martin, 1990; Petit et al., 1999).Variations in iron fluxes were presumably aconsequence of the areal extent of terrestrialdeserts and wind vectors. It is hypothesized thatincreased fluxes of iron to the high-nutrient, low

chlorophyll Southern Ocean stimulated phyto-plankton photosynthesis and led to a drawdown ofatmospheric CO2. Model calculations suggest thatthe magnitude of this drawdown could have beencumulatively significant, and accounted for theobserved variations in atmospheric CO2 recordedin gases trapped in the ice cores. However, thesedimentary records reveal large glacial fluxes oforganic carbon in low- and mid-latitude regions;areas that are presumably nutrient impoverished.Was there another factor besides iron additionto high-nutrient, low-chlorophyll regions thatcontributed to a net export of carbon duringglacial times?

Given that N:P ratios in the ocean interior arelower than for the sinking flux, an increase inthe net delivery of fixed inorganic nitrogen to theocean would also potentially contribute to a netdrawdown of CO2. The Antarctic ice core recordssuggest that atmospheric CO2 declined from,290 mmol mol21 to ,190 mmol mol21 over aperiod of ,8 £ 104 yr between the interglacialand glacial maxima (Sigman and Boyle, 2000).Assuming a C:N ratio of about 6.5 by atoms forthe synthesis of new organic matter in the euphoticzone, a simple equilibrium, three-box modelcalculation suggests that 600 Pg of inorganiccarbon should have been fixed by marine photo-autotrophs to account for the change in atmos-pheric CO2. This amount of carbon isapproximately threefold greater than that releasedto the atmosphere from the cumulative combus-tion of fossil fuels since the beginning of theIndustrial Revolution. The calculated change inatmospheric CO2 would have required an additionof ,1.5 Tg fixed nitrogen per annum; this is,2% of the global mean value in the contem-porary ocean.

8.05.14 LINKING IRON TO N2 FIXATION

Iron also appears to exert a strong constraint onN2 fixation in the modern ocean. Nitrogen-fixingcyanobacteria, like most cyanobacteria, requirerelatively high concentrations of iron (Berman-Frank et al., 2001). The high-iron requirementscome about because these organisms generallyhave high requirements for this element in theirphotosynthetic apparatus (Fujita et al., 1990), aswell as for nitrogen fixation and electron carriersthat are critical for providing the reductants forCO2 and N2 fixation in vivo. Increased Aeolian fluxof iron to the oceans during glacial periods mayhave therefore not only stimulated the utilizationof nutrients in high-nutrient, low-chlorophyllregions, but also stimulated nitrogen fixation bycyanobacteria, and hence indirectly provided asignificant source of new nitrogen. Both effectswould have led to increased photosynthetic carbon

Linking Iron to N2 Fixation 207

fixation, and a net drawdown of atmospheric CO2

(Falkowski, 1997; Falkowski et al., 1998).

8.05.15 OTHER TRACE-ELEMENTCONTROLS ON NPP

Is iron the only limiting trace element in thesea? Prior to the evolution of oxygenic photosyn-thesis, the oceans contained high concentrations(,1 mM) of dissolved iron in the form of Fe2þ

and manganese (.1 mM) in the form of Mn2þ,but essentially no copper or molybdenum; theseand other elements would have been precipitatedas sulfides. Thus, both iron and manganesewere readily available to the early photoauto-trophs. The availability of these two elementspermitted the evolution of the photosyntheticapparatus and the oxygen-evolving system thatultimately became the genetic template for alloxygenic photoautotrophs (Blankenship, 1992).Indeed, the availability of these transition metals,which is largely determined by the oxidation stateof the environment, appears to account for theiruse in photosynthetic reactions. However, overgeological time, oxygenic photoautotrophs them-selves altered the redox state of the ocean, andhence the availability of the elements in thesoluble phase (Anbar and Knoll, 2002).

As photosynthetic oxygen evolution proceededin the Proterozoic oceans, singlet oxygen (1O2),peroxide (H2O2), superoxide anion radicals (O2

z2),and hydroxide radicals (zOH) were all formed asby-products (Kasting, 1990; Kasting et al., 1988).These oxygen derivatives can oxidize proteins andphotosynthetic pigments as well as cause damageto reaction centers (Asada, 1994). A range ofmolecules evolved to scavenge or quench thepotentially harmful oxygen by-products; theseinclude superoxide dismutase (which converts O2

z2

to O2 and H2O2), peroxidase (which reduces H2O2

to H2O by oxidizing an organic cosubstrate for theenzyme), and catalase (which converts 2H2O2 toH2O and O2). The oldest superoxide dismutasescontained iron and/or manganese, while theperoxidases and catalases contained iron (Asadaet al., 1980). These transition metals facilitate theelectron transfer reactions that are at the core ofthe respective enzyme activity, and their incor-poration into the proteins undoubtedly occurredbecause the metals were readily available(Williams and Frausto da Silva, 1996). As O2

production proceeded, the oxidation of Fe2þ,Mn2þ, and S22 eventually led to the virtualdepletion of these forms of the elements in theeuphotic zone of the oceans. The depletion ofthese elements had profound consequences on thesubsequent evolution of life. In the first instance, anumber of enzymes were selected that incorpo-rated alternative transition metals that were

available in the oxidized ocean. For example,a superoxide dismutase evolved in the greenalgae, and hence higher plants (and many non-photosynthetic eukaryotes) that utilized copper andzinc (Falkowski, 1997). Similar metal substitutionsoccurred in the photosynthetic apparatus and inmitochondrial electron transport chains.

The overall consequence of the co-evolution ofoxygenic photosynthesis and the redox state ofthe ocean is a relatively well-defined trace-element composition of the bulk phytoplankton.Analogous to Redfield’s relationship between themacronutrients, trace-element analyses of phyto-plankton reveals a relation for trace elementsnormalized to cell phosphorus of (C125N16P1S1.3

K1.7Mg0.56Ca0.4)1000Sr4.4Fe7.5Mn3.8Zn0.80Cu0.38

Co0.19Cd0.21Mo0.03. The relative composition oftransition trace elements in phytoplankton isreflected in that of black shales (Figure 7).Thus, while there is a strong correlation bet-ween N:P ratios in phytoplankton and seawater(Anderson and Sarmiento, 1994; Lenton andWatson, 2000), there is no parallel correlationfor trace elements (Whitfield, 2001; Morel andHudson, 1985) (Figure 7 inset). There are twounderlying explanations for the lack of a strongrelationship. First, low abundance cations thathave a valance state of two or higher are oftenassimilated with particles, whether or not they aremetabolically required. Hence, while zinc,manganese, and cobalt are depleted in surfacewaters and are used in metabolic cycles, mercury,lanthanum, and other rare-earth elements are alsodepleted and have no known biological function.The profiles of these elements are dictated, to firstorder, by particle fluxes and their ligands. Second,the absolute abundance of transition metals iscritically dependent on the solubility of the sourceminerals, which is, in turn, regulated by redoxchemistry. Most key metabolic pathways evolvedprior to the oxidation of Earth’s atmosphere andocean, and hence many of the transition metalsselected for catalyzing biological redox reactionsreflect their relative abundance under anoxic orsuboxic conditions. For example, although thecontemporary ocean is oxidized, no known metalcan substitute for manganese in the water splittingcomplex in the photosynthetic apparatus. Simi-larly, the photosynthetic machinery has maintain-ed a strict requirement for iron for over 2.8 Ga,and all nitrogenases require at least 16 ironatoms per enzyme complex. Some key biologi-cal processes do not have the flexibility tosubstitute trace elements based simply on theiravailability (Williams and Frausto da Silva,1996). Hence, while oxygenic photoautotrophsindirectly determine the distribution of the majortrace elements in the ocean interior, the distri-bution of these elements in the soluble phase doesnot reflect with the composition of the organisms.


On long timescales, seawater concentrations oftrace elements reflect the balance between theirsources and sinks. For trace elements, the sinkterm is tightly coupled with the extent of redoxconditions throughout the ocean, where periods ofextended reducing conditions result in greaterpartitioning of organic carbon into the deep oceanand sediments. This, in turn, leads to enhancedsequestering of redox-sensitive trace elements intosediments, thereby decreasing their seawaterconcentrations. These sedimentary rocks have ahigh content of marine fractions (i.e., organicmatter, apatite, biogenic silica and carbonates),and so are enriched by 1–2 orders of magnitude inseveral trace elements: zinc, copper, nickel, moly-bdenum, chromium, and vanadium (Piper, 1994).The positive correlation between trace-elementratios in phytoplankton and sediments is consist-ent with the notion that phytoplankton haveimprinted their activities on the lithosphere.

8.05.16 CONCLUDING REMARKS

The evolution of primary producers in theoceans profoundly changed the chemistry ofthe atmosphere, ocean, and lithosphere of Earth.The photosynthesis processes catalyzed byensemble of these organisms not only influencesthe six major light elements, but directly andindirectly affect every major soluble redox-sensitive trace element and transition metal onEarth’s surface. These processes continue toprovide, primarily through the utilization of solar

radiation, a disequilibrium in geochemical pro-cesses, such that Earth maintains an oxidizedatmosphere and ocean. This disequilibrium pre-vents atmospheric oxygen from being depleted,maintains a lowered atmospheric CO2 concen-tration, and simultaneously imprints on the oceaninterior and lithosphere elemental compositionthat reflect those of the bulk biological materialfrom which it is derived.

While primary producers in the ocean compriseonly ,1% of Earth’s biomass, their metabolic rateand biogeochemical impact rivals the much largerterrestrial ecosystem. On geological timescales,these organisms are the little engines that areessential to maintaining life as we know it onthis planet.

ACKNOWLEDGMENTS

The author’s research is supported by grantsfrom the US National Science Foundation, theNational Aeronautical and Space Administration,the US Department of Energy, and the USDepartment of Defense.

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8.05 Biogeochemistry of Primary Production in the Sea Biogeochemistry of Primary Production in the Sea P. G. Falkowski Rutgers University, New Brunswick, NJ, USA 8.05.1 INTRODUCTION