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Global Biogeochemical Cycles and thePhysical Climate System

byFred T. Mackenzie

Atmosphere

Ecosphere

Hydrosphere Lithosphere

University Corporation for Atmospheric ResearchNational Center for Atmospheric Research UCAR Office of Programs

Understanding Global Change: Earth Science and Human Impacts

Global Biogeochemical Cycles andthe Physical Climate System

by

Fred T. MackenzieSchool of Ocean and Earth Science and Technology

University of HawaiiNational Oceanic and Atmospheric Administration

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Global Biogeochemical Cycles and the Physical Climate Systemby Fred T. Mackenzie

An instructional module produced by the Global Change Instruction Program of the UniversityCorporation for Atmospheric Research with support from the National Science Foundation.

GCIP Staff Advisory Committee

Tom M.L. Wigley, Scientific Director Arthur FewNational Center for Atmospheric Research Rice University

Lucy Warner, Program Manager John FirorUniversity Corporation for Atmospheric Research National Center for Atmospheric Research

Carol Rasmussen, Editor William MoomawUniversity Corporation for Atmospheric Research Tufts University

Linda Carbone, Secretary Ellen Mosley-ThompsonUniversity Corporation for Atmospheric Research The Ohio State University

Jack RhotonEast Tennessee State University

John SnowUniversity of Oklahoma

©1999 by the University Corporation for Atmospheric Research. All rights reserved.

Any opinions, findings, conclusions, or recommentations expressed in this publication are those of theauthors and donot necessarily reflect the views of the National Science Foundation.

For more information on the Global Change Instruction Program, contact the UCAR Communicationsoffice, P.O. Box 3000, Boulder, CO 80307-3000. Phone: 303-497-8600; fax: 303-497-8610;[email protected] or [email protected]://home.ucar.edu/ucargen/education/gcmod/contents.html

Global Biogeochemical Cycles and the Physical Climate System

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A note on this series

This series has been designed by college professors to fill an urgent need for interdisciplinary materialson global change. These materials are aimed at undergraduate students not majoring in science. Themodular materials can be integrated into a number of existing courses—in earth science, biology,physics, astronomy, chemistry, meteorology, and the social sciences. They are written to capture theinterest of the student who has little grounding in math and technical aspects of science but whose intel-lectual curiosity is piqued by concern for the environment. For a complete list of materials contactUCAR Communications (see previous page).

Global ChangeInstruction Program

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Contents

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .vii

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .1

Chapter 1: Bigeochemical Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .3

Chapter 2: Biogeochemical Cycles and Climate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .11

Chapter 3: The Modern Coupled C-N-P-S-O System . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .21

Chapter 4: Carbon Cycles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .23

Chapter 5: The Important Nutrient Nitrogen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .29

Chapter 6: Phosphorus and Sulfur . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .36

Chapter 7: The Water Cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .41

Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .44

Study Questions and Answers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .45

Glossary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .60

Supplementary Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .68

References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .69


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Preface

Global environmental change is a subject ofconsiderable public and scientific interest today.Any discussion of change must involve the sub-stances that are transported in cycles about theearth’s surface—through its air, water, soil, rocks,ice, and living and dead organic matter. Thinkingabout these global biogeochemical cycles andtheir role in environmental change requires us tocross the usual boundaries between biology, ecol-ogy, oceanography, meteorology, chemistry, andgeology. Because of the impact of human activi-ties on the cycles, and consequently the climate,the subject also involves the effects and conse-quences of natural and human-induced changefor ecosystems, humans, and human infrastruc-tures. This leads the discussion into the fields ofsociology, economics, and political science. Sucha broad and interdisciplinary topic is difficult tocapture completely in a module of this size. Ihave made no attempt to do so but have concen-trated on the biogeochemical cycles of five of themajor elements important to life—carbon, nitro-gen, phosphorus, sulfur, and oxygen—and theirrole in climatic change.

Biogeochemistry is the discipline that linksvarious aspects of biology, geology, and chem-istry to investigate the surface environment ofthe earth. This environment, the ecosphere (seeFigure 1), encompasses the biosphere (living anddead organic matter) and parts of the other largesubdivisions (reservoirs) of the earth’s surface ofatmosphere (air), hydrosphere (water), shallowcrust (soils, sediments, and crustal rocks), andcryosphere (ice). In this module, I focus on therole biogeochemistry plays in regulating andinteracting with the climate system.

This module covers a great deal of material,much of which is interdisciplinary. This presentsa problem for the writer of the material, theteacher, and the student. Both teacher and stu-dent generally will have more knowledge in one

discipline than in another. Also, each disciplinehas a unique vocabulary. (Most of the interdisci-plinary vocabulary in this module is definedwithin the text or in the extensive glossary.)Furthermore, the language of chemical equationsis used to describe processes operating within theecosphere. Therefore, it may take some additionalwork and perhaps reference to basic texts inchemistry, ecology, meteorology, etc., to digestthe material of this module.

The text begins by introducing some impor-tant biogeochemical processes. This material isnot a laundry list of processes but a selection ofsuch processes as photosynthesis, weathering,and deposition of sediments in the ocean asexamples of the nature and variety of biogeo-chemical processes. The next subject is the histor-ical (geological) nature of environmental changeon the earth. Emphasis is on the biogeochemicalcycles of atmospheric carbon dioxide and oxygenthrough the past 600 million years of the historyof the earth. The major processes controllingthese cycles and their tie to climate are discussed.We will see that for much of this time, the planethas had a more equable climate than at present.

Finally, the text deals with parts of the mod-ern biogeochemical cycles of five of the mostimportant elements essential for life: carbon,nitrogen, phosphorus, sulfur, and oxygen. Theseelements, along with hydrogen and a suite ofnutrient trace elements, interact through theprocesses of photosynthesis and respirationand/or decay. Processes and feedbacks withinthe cycles are described in the context of thepotential for a global warming brought about byhuman activities that have changed the composi-tion of the atmosphere. Keep in mind that theapproach can be used to interpret the interactionbetween biogeochemical cycles and climaticchange of any nature—warming or cooling—andat various space and time scales.


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An extensive glossary, study questions andanswers, and a supplementary reading sectionconclude the module. The glossary is interdisci-plinary and should help in understanding thediverse material of the module. The study ques-tions are designed to enable students to reviewthe text, to integrate the material, and to expandtheir knowledge of the topics covered. Many ofthe questions require calculations using standardarithmetic. Mathematics is the foundation of sci-ence, and it is necessary for students to get theirfeet wet. The readings are broad in scope and ofa general nature.

I would like to thank John Firor, DaveSchimel, and especially Tom Wigley for theircomments on the initial draft of this module.Some of the material in this module comes fromresearch supported by the National ScienceFoundation and the National Oceanographic and

Atmospheric Administration. The final version ofthis module was written while I was a Fellow atthe Wissenschaftskolleg zu Berlin. I thank Prof.Dr. Wolf Lepenies, rector of the institute, for pro-viding space, facilities, and peace of mind toaccomplish the task. Many thanks to MichaelShibao for drafting and in so doing substantiallyimproving the original illustrations for this mod-ule. Finally, I am extremely indebted to CarolRasmussen of the University Corporation forAtmospheric Research for her critical and labori-ous editing. Without her, this module would nothave been completed.

Fred T. MackenzieSchool of Ocean and Earth

Science and TechnologyUniversity of HawaiiJune 1996


1

The global ecosphere is the thin film aroundthe earth where living things (the biosphere1 )interact with the atmosphere (air), hydrosphere(water), cryosphere (ice), and lithosphere (soils andshallowly buried rocks) in a complex systeminvolving biological, geological, and chemicalprocesses and cycles (Figure 1). This biogeochemi-cal system of spheres and processes is poweredmainly by energy from the sun.

The ecosphere is made up of individualecosystems, such as tropical forests, grasslands,tundra, coral reefs, and estuaries. Matter andenergy flow between and within these ecosys-tems in interconnected biogeochemical cycles.Gaseous chemical compounds are produced andconsumed in the ecosystems and exchanged

between them and the air. In the atmosphere,they may react to form other compounds beforereturning to the earth’s surface. Some of thesechemical species are greenhouse gases, like carbondioxide (CO2) and methane, which act in theatmosphere to warm the planet. Others, likedimethylsulfide gas, react with other atmosphericchemicals to form minute airborne particles(aerosols) that directly or indirectly help to cool theclimate.

The most common way of studying the glob-al movements of these chemicals is by mathemat-ical modeling of biogeochemical cycles at theearth’s surface. Modeling also allows scientists toestimate the effects of human activities on naturalbiogeochemical cycles. A model is simply a set of

Introduction

Figure 1. The ecosphere, our life support system, showing its relationship to the other important spheres of the surface system of the earth(after Christensen, 1991).

Atmosphere

Ecosphere

Hydrosphere Lithosphere

1 Terms in italics are defined in the glossary at the end of the text.


2

equations that describe some of the processesfound in the real world. Biogeochemical cyclingmodels generally include processes that movematerials and their rates of transfer among a lim-ited number of well-studied spheres of the earth.

Biogeochemical cycles, however, have certainproperties that are inherently difficult to describeand model. These include: (1) irreversibility, thatis, the system does not return to its exact previ-ous state if it goes through a disturbance; (2) tran-sitional phenomena, that is, the system tends toswitch from one state to another and another andyet others, and perhaps back again, rather thansimply moving from “before” to “after”; (3) evolu-tion, in which the system progressively changes

in a particular direction; and (4) processes thateither enhance the original perturbation to thesystem (positive feedback) or relieve the perturba-tion (negative feedback).

In Chapter 1 of this module, we shall first con-sider some examples of biogeochemical processes.In Chapters 2 and 3, we shall discuss how the bio-geochemical cycles interact with climate, both inprevious eras and at present. In Chapters 4–6, weshall discuss the present-day global biogeochemi-cal cycles of several elements that are importantbiologically and that interact with the climate sys-tem. The cycles are looked at in the context ofglobal warming from an enhanced greenhouseeffect.

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Innumerable biological, geological, and chemi-cal processes cycle elements throughout the eco-sphere. The few discussed in this section shouldgive the reader an idea of their variety and com-plexity. As an example, consider a group of organ-isms called the prokaryotes: the bacteria and blue-green algae. The processes that these organisms areinvolved with (summarized in Table 1) include:

• the capture of carbon dioxide from theatmosphere and its conversion to organicmatter (fixation of CO2)

• the release of CO2 back to the atmosphere(through respiration and decay)

• fermentation of sugar

• methane production and oxidation

• sulfur reduction and oxidation

• nitrogen fixation, nitrification, and denitrification.

This list is given only as an example; some ofthese processes will not be discussed in the text.

These prokaryotic processes may take placein a variety of ways, such as (1) autotrophy, inwhich the organisms convert inorganic carbon inthe environment to organic matter; (2) heterotro-phy, in which the products from the breakdownof organic compounds are used to make neworganic materials; and (3) mixotrophy, in whichboth inorganic and organic compounds are usedto make organic matter.

Biogeochemical Processes

Table 1. Biogeochemical reactions involving prokaryotes

Element Process Summary of partial Examples of organismschemical reactions involved in process

Carbon CO2 fixation CO2 + H2 ⇒ (CH2O)n + Photoautotrophs:A2 (A = O, S) cyanobacteria, purple and green

sulfur bacteriaChemoautotrophs:

sulfur and iron oxidizing bacteria

Methanogenesis COO- + H2 ⇒ CH4 Methanogenic bacteria

Methanotrophy CH4 + O2 ⇒ CO2 Methanotrophic bacteria

Fermentation (CH2O)n + O2 ⇒ CO2 Anaerobic heterotrophic bacteria

Respiration (CH2O)n + O2 ⇒ CO2 Aerobic heterotrophic bacteria

Sulfur Sulfur reduction SO4 + H2 ⇒ H2S Sulfur-reducing bacteria

Sulfur oxidation H2S ⇒ S0 Purple and green sulfur phototrophs

S0 + O2 ⇒ SO4 Sulfur oxidizing bacteria

Nitrogen N2 fixation N2 + H2 ⇒ NH4 Phototrophic bacteria, nitrogen-fixing heterotrophic bacteria

Nitrification NH4 + O2 ⇒ NO2, NO3 Nitrifying bacteria

Denitrification NO2, NO3 ⇒ N2O, N2 Denitrifying bacteria__________________________________________________________________After Stolz et al., 1989



4

Principles of chemical reactions

Atoms and elementsEvery object in the universe is composed of matter. Because matter can be converted to energy, it is essen-

tially a form of energy. Matter is composed of atoms, which are the smallest particles of an element that canexist either alone or in combination. An atom is also the smallest particle that can enter into a chemical reac-tion. Most atoms never change; they only combine with other atoms to make different substances. Radioactiveatoms, however, do change and eventually decay into stable, nonradioactive atoms.

Elements consist of atoms of the same kind and, when pure, cannot be decomposed by a chemical change.There are 106 known elements; 103 are listed in the periodic table (Figure 2). The elements most used com-mercially by people, in order of use, are carbon (C), in the form of coal, oil, and gas; sodium (Na), in table saltand other products; iron (Fe), used in the steel industry; and nitrogen (N), sulfur (S), potassium (P), and calci-um (Ca), all used in fertilizers or as soil conditioners for our food supply.

CompoundsWhen two or more atoms are bonded together in a definite proportion, a compound is formed. Examples of

compounds discussed in this text are water (H2O), carbon dioxide (CO2), salt (NaCl), and sugar (e.g., glucose,C6H12O6). (All of the compounds named in the text are listed in Table 2.) The numbers in these chemical for-mulas are the number of atoms of each substance in the compound. If only one atom of a substance is in thecompound, no number is given. The universe is composed of millions of these compounds, all created from theelements given in the periodic table. The smallest particle of a compound that can exist and exhibit the prop-erties of that compound is called a molecule.

A compound is a pure substance that can be decomposed by a chemical change. The atoms in the chemicalcompound may rearrange themselves, or they may separate from the compound to form different compounds.These changes and interactions among compounds are called chemical reactions.

Chemical equationsA chemical equation expresses a chemical reaction involving compounds or elements. The chemicals that

react together, called reactants, generally are shown on the left-hand side of the equation and the products onthe right-hand side. Consider the decay of plant material (represented by the chemical compound CH2O, a car-bohydrate), which requires the oxygen gas (the chemical compound O2) in the earth’s atmosphere. The sim-plest chemical equation representing this process is

CH2O + O2 ⇒ CO2 + H2O (1)

The arrow pointing right indicates that this process is irreversible; the plant material will be completely oxi-dized to CO2 and H2O in the presence of atmospheric oxygen. Other processes are highly reversible, and theseare usually represented by a double arrow. For example, the equilibrium between calcium carbonate and itsdissolved calcium and carbonate ions (atoms or molecules that have lost or gained electrons, with the numberlost or gained shown as a positive or negative superscript) is represented as

CaCO3 ⇔ Ca2+ + CO32- (2)

In chemical processes, matter cannot be created or destroyed. Thus, when a chemical equation is written,the total number of atoms of any particular element on the left-hand side of a chemical equation must bemade to equal the total number of atoms of that element on the right-hand side of the equation. This is theprocess of balancing a chemical equation. Balancing the equation expresses the fact that molecules usuallyreact in such a way as to bear simple, integral, numerical relationships to one another.


5

If these relationships are known, it is possible to calculate the masses of reactants and products by usingknown atomic and molecular weights. In chemical terms, the amount of a substance is expressed in moles. Onemole of a substance is the amount that contains as many elementary entities as there are atoms in 12 grams ofcarbon. This number is termed Avogadro’s constant, and its value is equal to 6.022 x 1023. In the chemicalequation given above for the equilibrium of CaCO3 and its dissolved chemical species, one mole of CaCO3 willdissolve in water to make one mole of Ca2+ and one mole of CO3

2-. In terms of mass, 100 grams of CaCO3 willreact to give 40 grams of Ca2+ and 60 grams of CO3

2-. If only 10 grams of CaCO3 were to dissolve, then thesame proportions of Ca2+ and CO3

2- would be present at the equilibrium: 4 and 6 grams, respectively.

Figure 2. Periodic table of the elements. Each box includes an element’s atomic number,chemical symbol, and atomic weight.


6

Table 2. Chemical formulas and names used in this module

Al2Si2O5(OH)4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .kaoliniteCa2+ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .calcium ionCaCO3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .calcium carbonateCaSiO3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .calcium silicateCa5(PO4)3(OH,F) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .carbonate fluoroapatiteCH2O . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .carbohydrate(CH2O)106(NH3)16H3PO4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .organic matter in marine phytoplanktonCH4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .methaneCO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .carbon dioxideCO3

2- . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .carbonate ionCS2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .carbon disulfideC6H12O6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .sugar (glucose)DIC . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .dissolved inorganic carbonDMS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .dimethyl sulfide, (CH3)2SHCO3

- . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .bicarbonate ionHNO3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .nitric acidH2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .molecular hydrogenH2O . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .waterH2S . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .hydrogen sulfideH2SO4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .sulfuric acidH3PO4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .phosphoric acidH4SiO4

0 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .monomeric silicic acidKAlSi3O8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .orthoclase feldsparMSA . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .methane-sulfonic acidNaAlSi3O8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .albiteNaCl . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .sodium chloride, common table saltNH3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .ammoniaNH4

+ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .ammonium ionNH4NO3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .ammonium nitrate(NH4)2SO4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .ammonium sulfateNMHC . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .nonmethane hydrocarbonNO . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .nitric oxideNO3

- . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .nitrate ionNOx . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .oxides of nitrogenN2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .diatomic nitrogenN2O . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .nitrous oxideOCS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .carbonyl sulfideOH* . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .hydroxyl radicalOH- . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .hydroxyl ionO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .diatomic oxygen (pure oxygen molecules)O3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .ozonePAN . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .peroxylacetyl nitratePH3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .phosphine or swamp gasPO4

3+ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .phosphate ionSO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .sulfur dioxideSO4

2- . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .sulfate ionSOx . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .oxides of sulfurSiO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .silica


7DRAFT

Photosynthesis

We begin with perhaps the most importantbiogeochemical process of all, photosynthesis. It is aphotoautotrophic process, that is, an autotrophicreaction in the presence of light. Nutrients such asphosphate (PO4

3-) and nitrate (NO3-) are also nec-

essary for this reaction to occur. In the early stages of our planet’s formation,

the atmosphere was very different from that oftoday. There was no free molecular oxygen (O2),which most of today’s life forms require. In fact,oxygen was a very powerful poison for the sim-ple organisms that lived in this early, oxygen-deficient (anaerobic) world. Both the organismsand the earth had to evolve to a stage where theorganisms produced oxygen and emitted it totheir environment before more advanced lifeforms could evolve. Photosynthesis, the processof constructing complex organic molecules fromsimple inorganic ones in the presence of light,was a critical step in the evolution of life andallowed the mass of living organisms to grow tothe level of today. In our world, the mass of livingorganisms on earth is equivalent to about 600 bil-lion tons of carbon. More than 99% of this carbonis in land plants; the remainder is stored inmarine plants and in animals.

Photosynthesis is basically a chemical reactionor process in which carbon-, hydrogen-, and oxy-gen-bearing chemical compounds (carbohydrates)are synthesized from atmospheric CO2 and H2O oranother chemical compound that can act as ahydrogen donor. The generalized reaction is

energy + nCO2 + 2nH2A ⇒ (CH2O)n + nH2O + 2nA (3)

where H2A is a hydrogen donor molecule,(CH2O) is a carbohydrate, and n stands for anynumber. In higher plants, the donor molecule iswater, and n = 6. Thus for these plants the specificreaction is

energy + 6CO2 + 12H2O ⇒ C6H12O6 + 6H2O +6O2 (4)

For photosynthetic sulfur bacteria the donor mol-ecule is hydrogen sulfide (H2S), and for nonsulfurpurple bacteria it is organic compounds.

Incidentally, the carbon dioxide and donormolecule used for photosynthesis are not theonly requirements for plant growth. Plants alsoneed nitrogen, phosphorus, sulfur, potassium,and a dozen or so trace elements, like zinc andiron. As we shall see below, human activities arechanging the atmospheric concentrations of thesenutrients as well as that of carbon, with variouspossible effects on plants.

The photosynthetic reactions that produceorganic matter on land differ from those in theocean because the proportions of carbon, nitro-gen, sulfur, and phosphorus in land vegetationdiffer from those in marine plankton. The ratio ofC:N:S:P in marine plankton is 106:16:1.7:1.Known as the Redfield ratio, this proportion isfairly constant for the surface-dwelling, micro-scopic plants (phytoplankton) of the world’soceans. The C:N:S:P ratio for land plants is morevariable but averages 882:9:0.6:1. The amount ofcarbon is so much greater in land vegetationbecause it is stored as cellulose in the structuraltissues of trees and grasses.

From this summary, it can be seen thatphotosynthesis (among other things) links thebiogeochemical processes and cycles of the indi-vidual organic elements of carbon, nitrogen,phosphorus, and sulfur. These elements, plushydrogen and oxygen, are the major constituentsof organic matter. Those six elements and about adozen or so minor elements are necessary for themaintenance of organic structures and the physi-ological functions of living organisms.

Respiration and Decay of Organic Matter

In the life cycle, photosynthesis in plants isbalanced by the complementary processes of res-piration and decay in plants and animals. Inplants, respiration is the breakdown of the com-plex organic molecules that were formed duringphotosynthesis. The chemical reactions for respi-ration and decay are the reverse of those shownabove for the production of organic material. Thegeneralized reaction is

C6H12O6 + 6H2O + 6O2 ⇒ 6CO2 + 12H2O + energy (5)


8

Compare this with reaction 4. The amount ofenergy released is about 686 kilocalories (kcal) foreach mole of C6H12O6 (glucose, the most commonform of sugar in living things) that is broken down.

In animals, the respiratory oxidation of foods—that is, the loss of electrons from the carbon incarbohydrates, occurring during digestion—provides energy for a variety of uses, includingmaintenance of body temperature, muscularmovement, and synthesis of complex organiccompounds.

During the oxidation of organic matter, CO2,nitrogen- and phosphorus-bearing nutrients, andbioessential trace elements (e.g., iron) are returnedto the environment to be used again in the pro-duction of more organic matter. When O2 is avail-able, it is the oxidizing agent (oxidant); however,in oxygen-depleted (anoxic) waters, sediments,and soils, other oxidants are used. These includenitrate, sulfate, and iron and manganese oxides.

The chemical equations for respiration anddecay, either in an oxygenated or in an anoxicenvironment, are more complex than the general-ized reaction for photosynthesis given above. Forexample, the chemical composition of averagemarine phytoplankton—a relatively simple formof life—is (CH2O)106(NH3)16H3PO4: 106 mole-cules of carbohydrate, 16 of ammonia, and 1 ofphosphoric acid. When dead phytoplankton reactwith O2 in an oxygenated environment, the prod-ucts are carbon dioxide, nitric acid, phosphoricacid, and water:

(CH2O)106(NH3)16H3PO4 + 138O2 ⇒106CO2 + 16HNO3 + H3PO4 + 122H2O (6)

For an example of respiration and decay in ananoxic environment, let us consider the reductionof sulfur in sulfate (SO4

2-) in the pore waters ofanoxic sediments. Bacteria use the oxygen origi-nally bound in the sulfate to oxidize organic mat-ter. Again using phytoplankton as the organicmatter, the equation for this chemical reaction is

(CH2O)106(NH3)16H3PO4 + 53SO42- ⇒

106CO2 + 16NH3 + H3PO4 + 53S2- + 106H2O (7)

This time, in addition to carbon dioxide, phos-phoric acid, and water as in reaction 6, the prod-ucts include ammonia (NH3) and sulfide (S2-).

Weathering of Rocks

Another very important set of biogeochemi-cal processes is that involved with the break-down of rocks exposed to rain, wind, and ice.Weathering prepares rock for erosion and trans-portation. Its products are dissolved chemicalspecies and solids derived from changes in theprimary minerals of the rock being weathered.The solid products are predominantly clay min-erals; there are also dissolved products, predomi-nantly calcium, carbon, and silicon. Ultimately,the products of weathering are either carried bywater, blown as dust, or carried by glaciers to theocean. Of the approximately 20 billion tons ofsolids and dissolved materials reaching the oceanannually from the land, more than 80% is deliv-ered by rivers. However, high-temperature chem-ical reactions in the presence of seawater alongthe great submarine midocean ridges are signifi-cant sources of dissolved calcium, silica, and ironfor the oceans.

An example of a chemical weathering reac-tion is the weathering of the mineral albite (theinorganic chemical compound NaAlSi3O8), foundin igneous rocks like basalt, to the clay mineralkaolinite [Al2Si2O5(OH)4]. The reaction takesplace principally in the presence of soil waterand groundwater that contain significant amountsof dissolved CO2. Although the ultimate sourceof the CO2 is the atmosphere, much of it does notcome directly from the air but is produced insoils by the respiration of plants and the decay ofdead plants and animals. Because of theseprocesses, the concentration of CO2 in soils maybe one or more orders of magnitude greater thanthat of the atmosphere. The elevated CO2 levelsgive rise to acidic soil solutions, and these corro-sive, low-pH soil solutions are responsible for theweathering of rock minerals like albite:

2NaAlSi3O8 + 2CO2 + 11H2O ⇒Al2Si2O5(OH)4 + 2Na+ + 2HCO3

- + 4H4SiO40 (8)

The products of this reaction, besides thekaolinite, are sodium ion, bicarbonate ion, andmonomeric silicic acid.

In regions where human activities such ascoal burning release considerable amounts of sul-fur and nitrogen oxide gases to the atmosphere,


9

such as the midwestern and eastern United Statesand southern China, the pH of rainwater andconsequently soil water may be lower (moreacid) than natural values. This happens becausethe gases oxidize and react with water in theatmosphere and then rain out as sulfuric andnitric acids, respectively. This phenomenon is theenvironmental problem of acid deposition (oftencalled acid rain), which in extreme forms can beresponsible for increased fish mortalities in lakesand decreased agricultural production.

Deposition in the Oceans

When the solid and dissolved products ofweathering reach the ocean, the solids settle outbecause of their weight and are deposited on theseafloor as gravel, sand, silt, and mud. How longthe dissolved products remain in the oceandepends on how long it takes them to enter into achemical or biochemical reaction. As an exampleof the periods involved, dissolved sodium in theocean has a long residence time, about 55 millionyears. At the other end of the time scale, the resi-dence time of dissolved silica is only 20,000 years.

Calcium and silica

Many of the processes by which dissolvedconstituents are removed from the ocean involvemarine organisms. In today’s oceans, dissolvedcalcium and bicarbonate are precipitated as car-bonate minerals in the skeletons of several kindsof marine organisms: planktonic foraminifera (pro-tozoans), pteropods (mollusks), and Cocco-lithophoridae (algae), and bottom-dwelling (benthic)corals, echinoids, mollusks, and coralline algae. Ofthe total production of skeletal carbonate in theoceans, equivalent to about 1 billion tons of carbonper year, 80% is redissolved in the ocean as skele-tal debris sinks to the seafloor. This efficient recy-cling is due to the fact that although the surfaceocean is oversaturated with respect to calcium car-bonate, the deeper sea is undersaturated withrespect to this mineral. The remaining 20% of theocean’s carbonate production accumulates in shal-low-water and deep-sea sediments.

The amount of carbon in these sedimentsonly accounts for about one-half of the dissolved

inorganic carbon brought to the oceans annuallyby rivers. The other half of the riverborne carbonis released to the ocean and atmosphere whenskeletal carbonate minerals are formed.

Dissolved silica is also removed from theoceans in the skeletons of marine organisms.Certain of these organisms—planktonic diatoms(algae), radiolarians (protozoans), dinoflaggelates(protozoans), and benthic sponges—use dis-solved silica to form their shells of opaline silica.After these organisms die, most of the opal dis-solves, because the oceans throughout theirextent are undersaturated with respect to thischemical compound. Only about 40% of the totalannual production of skeletal silica sinks belowthe parts of the ocean that daylight reaches (theeuphotic zone). Most of this siliceous material dis-solves en route to the seafloor; only 5% of thatproduced in the euphotic zone accumulates inmarine sediments. This amount is about equiva-lent to the annual input of dissolved silica to theoceans by rivers.

Sodium and magnesium

In contrast to carbon and silica, which areremoved from the ocean primarily by biologicalprocesses, riverborne dissolved sodium and mag-nesium are removed to a significant extent byinorganic chemical reactions. Both of these ele-ments are involved in hydrothermal reactionsbetween seawater circulating through midoceanridges and the basalt rock making up the ridges.In the hydrothermal reaction process, sodiumand magnesium are removed from the seawater.Sodium is also removed from the ocean by theprecipitation of halite (common table salt, sodiumchloride) from seawater. This process is veryimportant as a removal mechanism for sodiumand chlorine, but only occurs when the right setof climatic and tectonic conditions are achieved.Only seawater in relatively isolated arms of thesea can be sufficiently evaporated to reach halitesaturation. Thus, because such environments arescarce today, it is likely that sodium and chlorinebrought to the oceans by rivers are currentlyaccumulating in seawater.

Some magnesium is also removed from sea-water by chemical processes in the pore waters of


10

sediments. These processes taking place duringthe burial of sediments are collectively referred toas diagenesis. The relative importance of diagenet-ic and hydrothermal reactions for the removal ofmagnesium from seawater is a topic of currentscientific research and debate.

We can conclude from the above discussionthat the circulation of material through the eco-sphere is complex and involves myriad chemical,biological, and geological processes. The system

is truly biogeochemical in nature. On all time andspace scales, if the composition of the ecosphereis regulated, the regulation is controlled by acomplex of interwoven inorganic and organicprocesses. The maintenance of the equable envi-ronment, including climate, that is required forlife to exist on earth is a product of this interact-ing and interwoven web of biogeochemicalprocesses and cycles.

11

In this chapter, we will look at some repre-sentative global biogeochemical cycles and theirrole in climate. The elements whose cycles arediscussed are intimately connected through theorganic processes of photosynthesis and respira-tion and/or decay. These elements are carbon(C), nitrogen (N), sulfur (S), oxygen (O), and—forcompleteness, because it is an important biologi-cal nutrient—phosphorus (P). We will begin witha discussion of gases whose production or con-sumption on the earth’s surface is accomplishedby biological reactions (biogenic gases) and thepossible effects of these gases on the earth’s cli-mate. To set the stage, the greenhouse effect isdiscussed briefly here.

Greenhouse gases, which are all naturallybiogenic in origin, allow incoming shortwave solarradiation to pass through the atmosphere to theearth’s surface, but when part of that radiation(about 45%) is reradiated back toward space asheat (infrared radiation), the gases absorb it andthus retain it in the atmosphere. This is the green-house effect. We can thank the natural green-house effect for the earth’s equable climate.Without it, the planet would be about 33°C coolerthan its mean global temperature of 15°C, that is,–18°C. (See the Global Change InstructionModule The Sun-Earth System, by John Streete.)

There can be, however, too much of a goodthing. In recent years, the concentrations of thesegases in the atmosphere have been rising becauseof fossil fuel combustion, biomass burning, ricepaddy cultivation, and other human activities.This buildup may absorb increased amounts ofoutgoing infrared radiation, leading to anenhanced greenhouse effect and global warming.

It is interesting and informative to put thispresent-day worry in the context of public andscientific concern about climate during the 1950sand 1960s. Between about 1940 and 1970, globalmean temperatures remained nearly constant, or

even declined slightly. There was considerablediscussion and concern in the scientific literatureand in public forums about global cooling andperhaps another ice age. Much of the discussionbelow in the context of global warming is applic-able to a scenario of global cooling as well. Thedifference is that in a global cooling, many of thefeedbacks mentioned would act in the oppositedirection and would probably have differentmagnitudes of change.

The Biogenic Gases and Climate

It is very likely that during the next centurythe earth’s climate will change due to naturalcauses. Changes in the amount of solar radiationreceived by the planet, in the circulation of theatmosphere and the oceans, and in volcanism canaffect climate on this time scale. On the longertime scale, if left to its own recourse, the planetwill most likely enter another ice age about10,000–30,000 years from now.

On the other hand, human-induced climatechange during the next century is also very like-ly. The flywheel of population growth and fossilfuel burning is turning rapidly and will be diffi-cult to slow in this time. The global population isgrowing at a rate of 1.5% per year, a doublingtime of 45 years. This rate of growth implies apopulation of about 10 billion by 2050. All thesepeople will require energy to sustain themselvesand to develop their industries, farms, and cities.Most scenarios of future global energy use pro-ject a continuous heavy reliance on fossil fuelinto the 21st century. Continued fossil fuel burn-ing will result in continued emissions of CO2,methane (CH4), and nitrous oxide (N2O) to theatmosphere. These greenhouse gases will beaccompanied by emissions of trace metals, non-methane hydrocarbons (NMHCs), oxides of sul-fur (SOx), and the most reactive oxides of nitro-gen (NO and NO2, collectively known as NOx).

Biogeochemical Cycles and Climate



12

The latter three groups of chemical compoundsreact with other chemical components of the cli-mate system, particularly the hydroxyl radical(OH*). Also, SOx and NOx are the principal con-stituents in acid deposition, and NOx andNMHCs are involved in the formation of ozone(O3), another greenhouse gas, in the troposphere.The unchecked accumulation of these gases inthe atmosphere could lead to an uncomfortablywarm planet.

The burning of fossil fuels and the burning offorests and other biomass are the principalhuman-induced, or anthropogenic, emissions ofmost biogenic gases to the earth’s atmosphere.Also, fossil fuel burning and changes in land use(such as deforestation) are responsible for manyof the global environmental problems the peopleof the world face today. Fossil fuel burning aloneaccounts for perhaps 80% of sulfur dioxide (SO2)emissions from the land surface to the atmo-sphere, 50% of carbon monoxide, 50% of NOx,20% of methane, 20% of NMHCs, 5% of ammo-nia, and 4% of nitrous oxide. It is also responsiblefor 70–90% of anthropogenic CO2 emissions tothe atmosphere. This amount is equivalent toabout 10% of the natural CO2 emissions from res-piration and decay.

As mentioned previously, C, N, P, and S,besides O and hydrogen (H), are the principalelements that make up living matter. The biogeo-chemical cycles of these elements are intimatelycoupled through biological productivity and respira-tion and/or decay. Ecosystems take energy fromtheir surrounding environment. The net result isproduction of organic matter, more disorder onthe planet (increased entropy), and waste. Thewaste may act as a pollutant. The biogenic gasesof carbon, nitrogen, and sulfur are a consequenceof this entropy production. Their fluxes maintainthe earth’s atmosphere in a state of disequilibrium.

The natural sources of these biogenic gasesare processes at the earth’s surface or chemicalreactions in the atmosphere. The processes bywhich biogenic gases and other componentscycle through the coupled C-N-P-S-O system,although in some environments operating closeto equilibrium, are principally controlled by therates at which the processes operate.

During the last half century, scientists havetended to specialize. Consequently, most globalenvironmental systems have been little studiedor studied only in a piecemeal fashion. Onlyrecently has attention been paid to the coupledearth-surface system of atmosphere, hydro-sphere, biosphere, cryosphere, and shallowlithosphere. Basic information on global reservoirsizes and fluxes (e.g., biological productivity) islacking or is only partly known. An example ofthis lack of data is the estimates of tropical forestbiomass, which vary by a factor of 2 or 3 forAmazonia alone.

It is very unlikely that the anthropogenicfluxes of gases to the atmosphere will substan-tially decline as we enter the 21st century.Population growth and our global reliance onfossil fuels as an energy source make such a sce-nario highly improbable. Thus, continued globalenvironmental change is a virtual inevitability. Itis likely that, by the middle of the next century,the atmospheric concentration of CO2 will bedouble what it was before the IndustrialRevolution (to date, it has increased about 30%),and concentrations of other greenhouse gaseswill also increase. Such a change in the composi-tion of the atmosphere portends a strong proba-bility of climate change.

Historical Framework

It is worthwhile considering at this stagehow the earth’s biogeochemical cycles and cli-mate system functioned prior to human interfer-ence. It is impossible to consider the functioningof all the biogeochemical cycles of concernbecause of space limitations. Only the global bio-geochemical cycles of carbon and oxygen areused as examples in this section. We will endwith a brief summary of environmental condi-tions just prior to major human interference inthe biogeochemical cycles and climate system.

Carbon

Carbon composes approximately 50% of allliving tissues. In the form of carbon dioxide, it isnecessary for plants to grow. Carbon dioxide alsohelps to sustain an equable climate on earth. Theconcentration of carbon dioxide in the


13

atmosphere has varied during the geologic past,but has remained within limits that permit life toexist on earth. Carbon dioxide is cycled through-out the spheres of earth on different time scales.We can refer to these scales as short-, medium-,and long-term. Figures 3 and 4 illustrate theprocesses involved in these time scales.

The short-term carbon cyclePhotosynthesis is part of the short-term car-

bon cycle (on the order of years). We can look atthe short-term cycling of carbon as carbon diox-ide by beginning with the producers of organiccarbon, the plants. Carbon in the form of atmo-spheric carbon dioxide is removed from the airby plants. This removal occurs both on land—forexample, in forests and grasslands—and inwater—for example, in lakes, rivers, and the

surface waters of the oceans. The primary pro-ducers, the photosynthetic phytoplankton andbenthic plants in the oceans and plants on theterrestrial surface, transform inorganic carbon ascarbon dioxide into organic carbon within theirtissues. Light and nutrients, like phosphate andnitrate, are necessary for this reaction to occur.Some of the energy from the light is used in thegrowth of plants, and some remains stored in thetissues of plants as carbohydrates.

Plants remove about 100 billion tons of car-bon as carbon dioxide from the global atmo-sphere each year, which is about 14% of theatmosphere’s total carbon. Most, but not all, ofthe carbon dioxide taken from the atmosphereduring photosynthesis is returned to the atmo-sphere during respiration and decay. The annual

(b) Subduction

(a) Photosynthesis-respiration

Uptake byrocks inweathering

(c) Stored oil, gas, coal,and kerogen

Organisms usecarbon from theocean/atmosphereto constructorganic matter andshells of calciumand carbonate,CaCO3

CO2

CO2

Decay

Figure 3. The biogeochemical cycle of carbon prior to human interference, showing (a) the short-term cycle, i.e., photosynthesis andrespiration; (b) the long-term cycle, involving accumulation of organic C and CaCO3 in marine sediments, their subduction, their alteration,and the return of CO2 to the atmosphere via volcanism; and (c) the medium-term cycle, involving storage of C in organic materials in sedi-mentary rocks. Ultimately this carbon is returned to the earth’s surface and undergoes weathering; in the process, O2 is taken out of theatmosphere and CO2 is returned.


14

removal rate of atmospheric carbon dioxide inphotosynthesis is slightly larger on land than inthe ocean.

After photosynthesis, carbon may next betransferred to a consumer organism if the plant iseaten for food. The carbon stored in the tissue of

the plant enters an animal’s body and is used asenergy or stored for growth. Land animals, suchas cows and deer, are the primary consumerorganisms. Aquatic plants are eaten by zooplank-ton (small sea animals) and larger animals. Whenan animal breathes, some of this carbon that was

ATMOSPHERECarbon in CO2 gas

CO2 fromdeforestation

CO2 fromdecaying

organic matterand

respiration

Oceansabsorb CO2

BIOSPHEREOrganic matter in

plants and animals

Living plantsextract CO2

CO2from

cementmanufacturing

CO2 toweatheringof limestoneand silicate

CO2from

burningof coal,oil, and

gas

HYDROSPHERECO2 dissolved in ocean

Aquatic plantsput CO2 into water

Buriedorganicmatter

Precipitationof CaCO3

Kerogen

Coal, oil, gas LITHOSPHERECarbon in buried plants, animals, and sediments

CO2 fromalteration of

organicmatter and

CaCO3

Figure 4. The major reservoirs and fluxes in the biogeochemical cycle of carbon. The shapes surrounding the spheres are called boxes.Arrows represent the processes and their directions that transfer carbon from one box to another. The carbon cycle can be conceived of asa series of interlocking circuits in the reservoirs of atmosphere, biosphere, hydrosphere, and shallow lithosphere (crust). In our time, thecycle would be in balance if it were not for human interference by burning of fossil fuels, cement manufacturing, and land-use activities(e.g., deforestation) (after Skinner and Porter, 1987).


15

taken up from plants is released from the ani-mal’s body as carbon dioxide gas.

Besides the carbon stored above ground inliving and dead vegetation, there is carbon belowground in the root systems of terrestrial plants.When the plants die, some of this carbon may bereleased as carbon dioxide or methane gas to theair trapped in the soil, or it may accumulate inthe soil itself as dead organic material. This deadorganic matter may be ingested by consumerorganisms, such as insects and worms living inthe soil.

Some of the organic carbon generated in landenvironments is weathered and eroded, and theorganic debris is transported by streams to theocean. In the ocean, some of this debris, alongwith the organic detritus of dead marine plantsand animals, settles to the ocean floor and accu-mulates in the sediments. However, some of thedebris is respired in the ocean to carbon dioxide.This carbon dioxide may leave the ocean and betransported over the continents, where it is usedagain in the production of land plants.

The long-term carbon cycleThe long-term carbon cycle (on the

order of tens to a hundred millionyears) requires that we consider theearth’s history over the last 600 millionyears or so—the period covered by thefossil record. Figure 5 defines theterms and intervals of geologic time.

The long-term cycling of carbon(Figures 3 and 4) involves intercon-nections between the cycling of theminerals calcium carbonate (CaCO3)and calcium silicate (CaSiO3). Thisseries of processes dates back to thebeginning of plate tectonics. Thiscycling includes not only the land andocean reservoirs but also that of lime-stone rocks. Limestone rocks are main-ly composed of calcium carbonate andare the fossilized skeletal remains ofmarine organisms or, less commonly,inorganic chemical precipitates of cal-cium carbonate. Limestones are greatstorage containers for carbon. Most ofthe carbon near the earth’s surface is

found in these rocks or in fossil organic matter insedimentary rocks. Weathering and erosion of theearth’s surface result in the leaching of dissolvedcalcium, carbon, and silica (SiO2) from limestonesand rocks containing calcium silicate.

The dissolved substances produced byweathering are transported to the ocean byrivers. As discussed in on p. 9, they are then usedto form the inorganic skeletons of benthic organ-isms and plankton, which are composed of calci-um carbonate and silica. During formation of thecalcium carbonate skeletons, the carbon dioxidederived from the weathering of limestone isreturned to the atmosphere.

When marine animals and plants die, theirremains settle toward the seafloor, taking the car-bon stored in their bodies with them. En route,their organic matter is decomposed by bacteria,just as on land. Some shells may dissolve. Thus,animal and plant organic and skeletal matter isturned back into dissolved carbon dioxide, nutri-ents, calcium, and silica in the ocean. This carbondioxide is stored in the deeper waters of the

Figure 5. The geologic time scale—the calendar of the earth. Geologic time isdivided into the intervals of eon, era, period, and epoch. The boundaries of theseintervals are based on absolute age dating using the radioactive decay of certainelements (e.g., uranium, potassium, rubidium, and carbon) in rocks; the distribu-tion of fossilized plants and animals found in the rocks; and certain worldwidegeologic events recorded in the rocks (after Skinner and Porter, 1987).

Eon Era Period Epoch Millions of years ago

Phanerozoic Cenozoic Quaternary Holocene

Pleistocene

Tertiary Pliocene

Miocene

Oligocene

Eocene

Paleocene

Mesozoic Cretaceous

Jurassic

Triassic

Paleozoic Permian

Carboniferous

Devonian

Silurian

Ordovician

Cambrian

Precambrian:

• Proterozoic

• Archean

Hadean

Today

0.01 (10,000 years ago)

1.6

5.3

23.7

36.6

57.8

65.0

144

208

245

286

360

408

438

505

545

2500

~3800

4600


16

oceans for hundreds to a thousand or so yearsbefore being returned to the atmosphere whenthe deep water moves upward (upwelling), usual-ly because of divergent movements of surfacewater.

Some of the animal and plant plankton sinksto the bottom, where the carbon in the organicmatter and shells escapes degradation andbecomes part of the sediment. As the seafloorspreads through plate tectonics, the sedimentscontaining the remains of marine plants and ani-mals are carried along to subduction zones, wherethey are transported down into the earth’s mantle.At the severe pressures and high temperatures inthe subduction zones, organic matter is decom-posed and calcium carbonate reacts with the sili-ca found in the subducted rocks to form rockscontaining calcium silicate.

During this metamorphism, carbon dioxide isreleased and makes its way into the atmospherein volcanic eruptions and via hot-spring dis-charges. Once in the atmosphere, it can thencombine with rainwater. The rainwater falls onthe land surface and seeps down into the soils,where it picks up more carbon dioxide fromdecaying vegetation. This water, enriched in car-bon dioxide, weathers and dissolves the com-pounds of calcium and silica found in rocks ofthe continents. The cycle begins again.

This series of processes has been active for atleast 600 million years, since the advent of thefirst organisms that made shells (and were there-fore the first to leave fossils). The processes wereimportant even earlier in earth’s history, whencalcium carbonate was deposited in the ocean byinorganic processes.

Figure 6. Model calculation of atmospheric carbon dioxide during the last 600 million years. The horizontal axis shows time in millions of yearsbefore the present (top) and geological time period (bottom). The left vertical axis shows the number of times today’s C level that existed in theatmosphere of the time; the right vertical axis shows the amount of CO2 in the atmosphere. For example, 500 million years ago there wasabout 14 times as much CO2 in the atmosphere as there is today, with a total amount of about 37 x 1018 grams (after Berner, 1991).

KTr

18

16

14

12

10

8

6

4

2

0

20

30

40

50

10

0

600 500 400 300 200 100 0

TimeMillions of years before present (BP)

CO

2 in

the

atm

osp

here

rel

ativ

e to

pre

sent

day

Atm

osp

heric CO

2 (1018 g

rams)

Paleozoic Mesozoic Cenozoic

C O S D C P J T

C - CambrianO - OrdovicianS - SilurianD - DevonianC - CarboniferousP - PermianTr - TriassicJ - JurassicK - CretaceousT - Tertiary

Present day


17

Atmospheric CO2 levels of the pastOne outcome of changes in the rates of

processes in the long-term biogeochemicalcycling of carbon is that atmospheric carbondioxide has varied in a quasicyclic fashion dur-ing the last 600 million years of earth’s history.Robert A. Berner of Yale University and col-leagues have developed models of the carboncycle to calculate these variations. Figure 6 showsthe results of one such calculation.

Periods of high atmospheric carbon dioxidelevels are the result mainly of intense plate tec-tonic activity, with increased metamorphism oflimestone and release of carbon dioxide to theatmosphere from volcanoes. These high carbondioxide periods are often referred to as hothousesor greenhouses. They are also periods of relative-ly high sea level; for example, in the Cretaceousand Cambrian periods, much of what wouldbecome the present continent of North Americawas covered by water. This flooding of the conti-nental landscapes was due principally to thelarge size and volume of the midocean ridges,caused by intense plate tectonic activity. Theincrease in ridge volume led to the displacementof ocean water onto the continents of the time.

Periods with lower atmospheric carbon diox-ide levels, such as the Carboniferous throughTriassic and much of the Cenozoic, are the out-come of less intense plate tectonic activity andincreased removal of carbon dioxide from theatmosphere by weathering. These intervals,which include the present era, are extended coolperiods (ice houses) in the climatic history of theearth. They are also times of relatively low sealevel due to a decrease in the volume of themidocean ridges.

Biological and other factors also play a rolein regulating atmospheric carbon dioxide levelsover the long term. For example, the lowering ofatmospheric CO2 from the high levels of the mid-Paleozoic era, about 400 million years ago, wasnot simply the result of the waning intensity ofplate tectonic processes. It followed the evolutionof land plants, which withdrew CO2 from theatmosphere by photosynthesis. Similarly, thelowering after the Cretaceous period, about 100million years ago, followed the appearance offlowering plants, which resulted in an increase in

weathering rates and withdrawal of CO2 fromthe atmosphere.

The important point is that atmospheric CO2has varied by a factor of perhaps more than tenduring the last 600 million years of the earth’shistory. This variation certainly has had climaticimplications, because CO2 is an important green-house gas. In fact, for much of the last 600 mil-lion years, the planet had a different atmosphericcomposition and a more equable climate thanthat of today.

The medium-term carbon cycleMedium-term cycling of carbon dioxide (mil-

lions to tens of millions of years) involves organicmatter in sediments; coal, oil, and gas; andatmospheric oxygen (Figures 3 and 4). It com-mences, as does the short-term cycling, with theremoval of carbon dioxide from the atmosphereby its incorporation into plants and the accumu-lation of the dead plant and animal carbon insedimentary organic matter (the dead and fos-silized remains of plants and animals). Whendispersed throughout a sedimentary rock, thisorganic matter is termed kerogen. Shales are veryfine-grained sedimentary rocks that are often richin kerogen. Coal, oil, and gas deposits are alsothe altered remains of the soft tissues of plantsand animals which have accumulated in a geo-graphically restricted area.

Coal is derived mainly from terrestrial plantmaterial, which is often deposited in swampyenvironments. The plant material is altered whenthe swamp sediments are buried. If buried deepenough, the dead plant material may be substan-tially changed because of the increased tempera-tures and pressures at depth. Different types ofcoal are formed by varying conditions of temper-ature and pressure. Anthracite—a hard, dense,black coal—is formed by alteration of plant mate-rial at a relatively high temperature and pressure.Bituminous, brown coal is formed under lessintense conditions. Peat, used as a fuel in someparts of the world, is little-altered plant materialthat has not been buried deeply. It is high in car-bon. Peat is an important component of tundraareas in the Northern Hemisphere.

Oil and gas represent highly altered organicmatter, principally the altered remains of marine


18

phytoplankton that were deposited on theseafloor. During burial, these organic materialsare broken down at elevated temperature andpressure, forming oil and gas. The oil and gasmay migrate hundreds of miles in the subsurfacebefore coming to rest in large accumulations inthe voids of rocks. Often, oil and gas are formedin shales; but as temperature and pressureincrease with depth, they move to more coarse-grained rocks like sandstones and limestones. Itis in these latter rocks, dating from the Creta-ceous and Cenozoic periods, that the great oiland gas reserves of the world are found, likethose of the Persian Gulf.

These deposits of coal, oil, and gas comefrom organic carbon that has escaped respirationand decay. Thus, they represent carbon dioxidethat has been removed from the atmosphere. Thesame is true of the kerogen dispersed as fine-grained materials in the sedimentary rocks.Because these materials were buried, the oxygenthat would have been used in their decay hasremained in the atmosphere. Eventually, howev-er, the carbon in these deposits and in kerogen isrecycled into the atmosphere, returning carbondioxide to that reservoir and removing oxygen.This happens when these fossil fuel deposits andkerogen are uplifted by plate tectonic forces aftermillions of years of burial and exposed to theatmosphere. When this occurs, the oxygen thatpreviously accumulated in the atmosphere reactswith the coal, oil, gas, and kerogen. The reactioninvolves the decay of these organic materials inthe presence of oxygen (equation 5). This resultsin the removal of oxygen from the atmosphereand the return of carbon dioxide to the atmo-sphere. The ongoing dynamic cycle is complete.

Fossil fuel is a nonrenewable energy source,because coal, oil, and gas deposits take millionsof years and specific environmental conditions toform. The mining of these deposits brings thesematerials back to the surface much more rapidlythan natural processes do. The stored energyfrom the long-dead organisms is released in theform of heat when the coal, oil, and gas areburned. This fossil fuel energy keeps us warm,powers our cars, and moves the machinery ofindustry. It is also a main cause of environmentalpollution, because a byproduct of fossil fuel

burning is the release of gases and particulatematerials into the environment. Climatic changeis an important potential global environmentaleffect of the release of carbon dioxide and othergases to the atmosphere by combustion.

In summary, carbon is found in all four majorsurface spheres of the earth. In the ecosphere, it isessential to every life form, occurring in allorganic matter. In the atmosphere, it is found asthe gases carbon dioxide, methane, and severalother compounds. It occurs as carbon dioxidedissolved in lakes, rivers, and oceans in thehydrosphere. In the earth’s crust—part of thelithosphere—it is found as calcium carbonateoriginally deposited on the seafloor, as kerogendispersed in rocks, and as deposits of coal, oil,and gas. It is because carbon is stored in the largesedimentary reservoirs of limestone and fossilorganic carbon, and not in the atmosphere, thatlife on earth is possible. If all this carbon werestored in the atmosphere, there would be about30 times the existing amount of CO2. The resultwould be a substantial heating of the atmospherebecause of the increased absorption of the long-wave infrared radiation emitted from the surfaceof the earth. The greenhouse effect would be verystrong, and our planet would probably have atemperature like that of Venus: 460°C!

Oxygen

Oxygen composes 20.9% of the gases of theatmosphere. The cycling of oxygen (Figure 7) isstrongly coupled to that of carbon (Figure 4).Oxygen is produced by plants during photosyn-thesis, when carbon dioxide is consumed. It isremoved by respiration and decay, when carbondioxide is produced. This is a short-term, nearlybalanced cycle on land, because the amount ofoxygen produced yearly by land plants is aboutequivalent to the amount they use in the processesof respiration and decay. On average, it takesabout a decade for oxygen and carbon dioxide tocycle through living plants. However, a little ter-restrial organic matter that has not undergonerespiration or decay, such as leaves and trunks oftrees and smaller-sized organic debris, leaks fromthe land to the ocean via rivers. In the ocean,some of this organic detritus escapes destructionand is deposited in the sediments.


19

In the oceans, phytoplankton produce slight-ly more oxygen than is consumed during the res-piration and decay of marine life. As a result,oxygen is released to the atmosphere. The organ-ic carbon not decayed by this oxygen, along withsome of the terrestrial organic detritus mentionedin the preceding paragraph, is deposited on theseafloor and accumulates in the sediments of theocean. If this accumulation were not counteractedby other processes, and no other factors were tointervene, all the carbon dioxide in the atmo-sphere would disappear in less than 10,000 years.The oxygen content of the atmosphere woulddouble in less than several million years.Fortunately, the overproduction of oxygen in theoceans is balanced by the weathering of fossilorganic carbon and other materials in rocks onland. During this process, carbon dioxide isreturned to the atmosphere.

If the burial of organic carbon in sedimentswere enhanced, oxygen might accumulate in the

atmosphere. In fact, it seems to be the case thattimes of high organic carbon burial in the pastgave rise to high atmospheric oxygen levels.Figure 8 is a model calculation of atmosphericoxygen concentration variations during the last600 million years. Just before the Carboniferous,vascular plants evolved and spread over the conti-nents. Their remains were a new source of organ-ic matter resistant to degradation by atmosphericoxygen. During the Carboniferous and Permian,large quantities of vascular plant organic matterwere buried in the vast coastal lowlands andswamps of the time. This material became thecoal deposits mined from rocks of Carboniferousand Permian age today. This large accumulationof organic matter gave rise to the high atmospher-ic oxygen levels of the late Paleozoic. Coaldeposits are also important in Cretaceous- andearly Cenozoic–age rocks, another time of highatmospheric oxygen concentrations.

Environmental Conditions BeforeHuman Interference

From the above discussion, we can concludethat the carbon and oxygen biogeochemicalcycles have changed throughout geologic time.These changes have led to changes in atmospher-ic composition and climate. Just prior to signifi-cant human interference in the ecosphere, envi-ronmental conditions were also changing natu-rally because of a variety of factors. These factorsinclude external forcings, such as changes in theorbit of the planet earth and in the flux of solarradiation to the planetary surface, and internalforcings, involving the behavior of the earth’socean-atmosphere-land-biota-cryosphere system.

For about 1.6 million years, the earth hasexperienced a series of cold and icy times knownas ice ages or glacial stages. These periods ofextensive glaciation of the continents alternatedwith shorter and warmer periods, interglacialstages like today. During the interglaciations, thegreat glaciers that covered the continents ofEurope and North America to depths of two kilo-meters or more melted. The glaciers retreated togeographical positions similar to those of the pre-sent continental ice sheets of Greenland andAntarctica. As the glaciers waxed and waned

Atmospheric oxygen

AtmosphericCO2

Oxygen dissolvedin the ocean

Land biota

Marine biota

RespirationPhotosynthesis

Phot

osyn

thes

is

Resp

iratio

n

Marine sediments(organic carbon)Continental rocks

Figure 7. The biogeochemical cycle of oxygen. This cycle is strong-ly coupled to that of carbon (Figure 4). The boxes represent themajor reservoirs of oxygen, and the arrows the fluxes of oxygenfrom one box to another. The heavier the arrow, the larger the flux.The broken lines show the flow of carbon in sedimentation on theocean floor, burial in sediments, and uplift by plate tectonicprocesses. When uplifted, organic C is oxidized by oxygen in theatmosphere, and CO2 is released (after Andreae, 1987).


20

through the last 2 million years, global tempera-tures went up and down, as did sea level. Thecomposition of the atmosphere and other envi-ronmental conditions changed.

The earth, before extensive human interfer-ence in its biogeochemical cycles, was recoveringfrom the climax of the last great glaciation 18,000years ago. The recovery has not been smooth.There have been times in the past 18,000 yearswhen the planet cooled quickly. Also, there havebeen periods when the earth was warmer than

today. However, during the past several cen-turies, the global environment has changed sub-stantially and rapidly. Atmospheric trace gas con-centrations, matter in runoff, temperature, andother indicators have increased in magnitude. Amajor reason for all the changes is the impact ofhuman activities on the environment.

10

0

20

30

35

10

0

600 500 400 300 200 100 0

TimeMillions of years before present (BP)

Oxy

gen

(1020

gra

ms)

O2 as p

ercentage o

f atmo

spheric g

ases

Paleozoic Mesozoic Cenozoic

C O S D C P Tr J K T

C - CambrianO - OrdovicianS - SilurianD - DevonianC - CarboniferousP - PermianTr - TriassicJ - JurassicK - CretaceousT - Tertiary

Present day

20

30

25

15

5

Figure 8. Model calculation of atmospheric oxygen during the past 600 million years. The dashed line across the figure shows today’satmospheric concentration of O2. The left vertical axis shows the atmospheric O2 level; the right vertical axis shows O2 as a percentage ofthe total atmospheric gases. The horizontal axis shows time in millions of years before the present and geological time period (after Bernerand Canfield, 1989).

21

The modern coupled C-N-P-S-O system ismultidimensional and complex, with numerousprocesses, reservoirs, and fluxes. All these attrib-utes are difficult to portray in detail. To see therelationships between biogeochemical cycles andclimate, we will examine those biogeochemicalcompounds that play a role in climatic change. Inparticular, we will focus on the problem of anenhanced greenhouse effect and global warming

brought about by human activities. Keep in mindthat the cycles also respond to a global coolingand can feed back into any such cooling. Theprocesses discussed below would respond tocooling in an opposite sense, to some extent, andprobably with different magnitudes of change.

Figures 9–20 show the biogeochemical cyclesof interest. The figures show that the burning offorests and other biomass and fossil fuel

The Modern Coupled C-N-P-S-O System

METHANE(fluxes = Mt C/y)

Foss

ilfu

elbu

rning

34– 76

Biom

ass b

urnin

g16

–31

Land

fills

16–

53Ri

cepa

ddie

s19

–12

7CH

4hy

drat

e0

–76

Soil

upta

ke11

–34

Productionfrom

live stock56

–150

Freshwater

1–

19

Natural wetland s

76–

150

Ocean

4–

16

Mainly OH*depletion

300 – 450CO

CH4 + OH* → H2O + CH3

To s

trat

osph

ere

46Climate sensitivity0.3°C/CH4 doubling

17% greenhouse

CH43750 Mt C1.72 ppmv

Residence time = 10 yAccumulation:

28 Mt C/y8 ppbv

Land Ocean

Figure 9. Part of the modern global biogeochemical cycle of methane, emphasizing exchanges of methane between the earth’s surface andatmosphere and the fate of the gas in the atmosphere. The atmospheric reservoir is shown as a circle, including the amount of carbon interagrams (1012 grams), equivalent to million metric tons (Mt). Flux ranges are also given in Mt C/yr; arrows indicate the directions of thefluxes. The proportion of C in the atmosphere is given in parts per million by volume (ppmv), the residence time in years, and the accumula-tion rate per year in both Mt and parts per billion by volume (ppbv). The chemical reaction at the top of the figure shows the fate of methanethat escapes to the stratosphere (after Mackenzie, 1995; Houghton et al., 1996).



22

combustion are the major human sources of mostbiogenic gases in the earth’s atmosphere. Manyof the natural exchanges of gases between theearth’s surface and the atmosphere in thesecycles are driven by biological processes, whichinvolve bacteria. In the atmosphere, many of thegases are oxidized, usually by hydroxyl radical(OH*), a trace gas that is the atmosphere’s maincleansing mechanism. The oxidized materialsthen return to the earth’s surface, through biolog-ical production of plants on land and in theocean and by wet and dry deposition.

Many of the natural processes shown in thesefigures have feedbacks that affect the accumula-tion of trace gases in the atmosphere and henceaffect the global climate. Because these feedbacksare linked to biological processes, they are termedbiotic feedbacks. We shall explore the nature ofthe feedbacks in the C-N-P-S-O system below.

CARBON MONOXIDE(fluxes = Mt C/y)

Foss

ilfu

elbu

rning

336

Biom

ass

burn

ing31

2

Soil

48

Ocean72To

str

atos

pher

e

48

OH* depletion CO2

1212

OH* depletion NMHC360

CO + OH* → CO2 + H

OH* depletion

CH4

324

Soil

upta

ke

192

CO228 Mt C

0.11 ppmvResidence time = 0.2 y

Accumulation:2 Mt C/y1 ppbv

Land Ocean

Figure 10. Part of the modern global biogeochemical cycle of carbon monoxide, emphasizing exchange of the gas between the earth’ssurface and its atmosphere and the fate of the gas in the atmosphere. See Figure 9 for an explanation of the units and abbreviationsused. Notice, as with methane, the role of OH* as an agent of oxidation of the reduced carbon gases (after Mackenzie, 1995; Houghtonet al., 1996).

23

The CH4-CO-CO2 Connection

Figures 9–11 illustrate the biogeochemicalcycles of methane (CH4), carbon monoxide (CO),and carbon dioxide between the earth’s surfaceand atmosphere. We have a reasonably goodunderstanding of the major processes associatedwith the exchange of these gases between thesurface and the atmosphere and of the gases’atmospheric reservoir sizes. However, the fluxesassociated with the various processes are lesswell quantified.

All three of these gases have been accumulat-ing in the atmosphere because of human activi-ties. Since the late 18th century, global atmo-spheric CO2 concentrations have risen fromabout 280 to about 360 parts per million by vol-ume (ppmv), a change of approximately 30%.This increase is principally the result of fossil fuelburning and land-use activities such as deforesta-tion, which add CO2 to the atmosphere when theburning carbon combines with atmospheric mol-ecular oxygen. Because of this increase in CO2from human activities, the atmospheric reservoirof oxygen gas has been reduced by less than 1%.Thus, in the burning of fossils fuels, there is noconcern about depletion of atmospheric CO2.

The global concentration of CH4 during thisperiod has more than doubled from about 800 to1,720 parts per billion by volume (ppbv). As withCO2, this increase is due to human activities, andagain as with CO2 those activities include burn-ing of biomass and fossil fuels. For methane,though, rice paddies, farm and ranch animals(which produce methane as they chew their cud),and rotting landfills are also important. Landfillsemit CH4 as methanogenic bacteria decompose thewastes in an anaerobic environment.

Carbon monoxide has also been increasing inconcentration in the atmosphere, particularly inthe Northern Hemisphere where important

anthropogenic emission sources are located. Therate of increase in the 1980s was approximately 1ppbv per year, principally because of the same cul-prits, fossil fuel and biomass burning (Figure 10).

The rates of increase of all three atmosphericcarbon gases slowed during the late 1980s andearly 1990s because of a variety of natural andhuman-induced causes.

Major processes in the CH4 and CO cycles

CH4 and CO are reduced gases, that is, theymay react with other chemical compounds by theloss of electrons from the carbon in the com-pounds. The carbon in CH4 has a valence of –4;the carbon in CO has a valence of +2. The biogeo-chemical cycles of these two gases and CO2 areconnected because OH* oxidizes CH4 and CO toCO2, which has a valence of +4. This CO2 canthen be used by plants and other organic matter.The decay of this organic matter, in turn, leads tothe release of CH4 and CO from the earth andoceans.

The major natural processes involvingexchange of CH4 between the earth’s surface andthe atmosphere (shown on the right-hand side ofFigure 9) are

• evasion (release) to the atmosphere from theocean and natural wetland and freshwaterecosystems,

• leakage from underground natural gasdeposits into the atmosphere (in Figure 9,this is included in the fossil fuel burningflux), and

• uptake of CH4 in soils due to the activity ofmethanotrophic bacteria.

The major sink of CH4 is the atmosphere,where it is oxidized principally by reaction withOH* (top right-hand side of Figure 9).Approximately 300–450 million tons of carbon

Carbon Cycles



24

annually are oxidized this way. During the pasthundred years, the change in the concentration ofmethane in the atmosphere could be responsiblefor about 20% of the warming due to anenhanced greenhouse effect. If CH4 concentrationwere to double in the atmosphere, one wouldexpect, based on the ability of the gas to warmthe atmosphere, a rise of approximately 0.3°C inglobal mean temperature.

Soils are both a natural source of CO to theatmosphere and a natural sink of the gas fromthe atmosphere (Figure 10). The source is the bac-terial decomposition of organic matter in soils,and bacteria and algae in the ocean. The sink isalso the result of bacterial processes.

The atmosphere is also both a source andsink of CO. It is the most important sink of CO,as with CH4, through the oxidation of CO to CO2via OH*. This process transfers approximately1,200 million tons of carbon to the atmosphericCO2 reservoir annually (Figure 10). The atmo-sphere is a source of CO because nonmethanehydrocarbons (NMHCs) are oxidized to thatcompound by OH*.

CH4 and CO feedbacks to global climate change

Respiration and the bacterial decompositionof organic matter emit CH4 and CO from soils tothe atmosphere. Because the rates of theseprocesses increase with increasing temperature,the emissions might increase in a world that waswarming. This would be a positive feedback onany initial warming of the earth. However,because soils are also sinks for atmospheric CH4and CO, and emissions of these gases are sensi-tive to the amount of moisture in the soil, the sit-uation is more complex.

In general, it is likely that warmer tempera-tures in the high northern latitudes will lead toan increase in CH4 fluxes from CH4 trapped inpermafrost, decomposable organic matter frozenin permafrost, and decomposing CH4 gashydrates—substances in which CH4 is locked in astructure of water ice. These hydrates are stableonly at low temperatures or under pressuresexceeding ten atmospheres. They are stored insediments under shallow seas, particularly in theArctic region, and also in permafrost. The flux of

methane escaping to the atmosphere from thissource is poorly known today (Figure 9).

Because CH4 and CO react easily with OH*in the atmosphere, increases in their atmosphericconcentrations (and in those of the reduced Nand S gases) would affect the concentration ofOH*, thus having an impact on the atmosphere’sability to cleanse itself. It is possible that anincrease in the flux of CO to the atmosphere,because of the gas’ short residence time there,could deplete OH* and consequently cause lessCH4 to be removed from the atmosphere. If so,CH4 would accumulate faster in the atmosphere,leading to an increased rate of global warming.However, the overall effect is likely to be small.Furthermore, increased atmospheric CH4 wouldprobably lead to increased production of watervapor in the stratosphere (Figure 9), wheremethane and OH* react to form water vapor andCH3. Water vapor is a greenhouse gas, and itsincreased production in the stratosphere is a pos-itive, but minor, feedback in a scenario of awarming earth.

In summary, although the various effects aredifficult to quantify, it is likely that an initialwarming of the climate would lead to a netincrease in fluxes of CH4 and CO to the atmo-sphere and accumulation there. This situationconstitutes a positive feedback on the concentra-tion of these gases in the atmosphere and henceon global warming. Any global cooling wouldprobably result in an opposite effect.

The CO2 Cycle

The most important trace gas contributing tothe potential of an enhanced greenhouse effect isCO2. From 1765 to 1995, CO2 accounted for about60% of the human-induced warming due to theaccumulation of greenhouse gases in the atmo-sphere. A doubling of the concentration of CO2 inthe atmosphere could eventually lead to anincrease in global mean temperature of about1.5–4.5°C.

Figure 11 shows the major processes, bothnatural and human, affecting the exchange ofCO2 between the atmosphere and earth’s surface.Estimates of the fluxes associated with theprocesses are also shown. Let us begin our


25

exploration of Figure 11 by considering the oceanfluxes. The total exchange of CO2 between theocean and atmosphere amounts to about 90 bil-lion tons of carbon per year. Much of this carbonexchange involves

• evasion of CO2 from the warm surfacewaters of the tropics and from upwellingregions, and

• invasion (uptake) of CO2 into higher-latitude,colder, surface waters (see equation 9, p. 27).

Almost half of this exchange—close to 45 bil-lion tons of carbon annually—is involved in netprimary production due to CO2 fixation by

photoautotrophic marine plankton. A slightlylarger amount of CO2 is returned to the atmos-phere by respiration and decay in the ocean. Thedifference between the two amounts implies thatthe ocean, prior to human interference in theglobal biogeochemical cycle of carbon, was a netheterotrophic system and supplied CO2 to theatmosphere. The CO2 released from the ocean wassubsequently used in organic production on land.

In contrast, the terrestrial biota were a netautotrophic system before humans interfered withthe carbon cycle. In other words, the land biotaproduced more organic carbon than was con-sumed. (Compare the rates for net primary

CARBON DIOXIDE(fluxes = Mt C/y)

Land

use

1600

Foss

ilfu

elbu

rning

6100

Resp

iratio

n-de

cay

6144

0Ne

t prim

ary

prod

uctio

n63

000

Wea

ther

ing

216

Volcanic

60

CaCO3 precipitation

168

Respiration-decay

45252

Net primary production

45000

OH* depletion

CO

Climate sensitivity2.5°C/CO2 doubling

60% greenhouse

Oxidationof fossil CH

2 O

36

1212

Excess4412

To ocean2000

To terrestrial realm2412

Late diagenesis-metamorphism60

CO2752000 Mt C

360 ppmvResidence time = 6.5 y

Accumulation:3100 Mt C/y

1.5 ppmv

Land Ocean

Figure 11. Part of the modern global biogeochemical cycle of carbon dioxide, emphasizing its exchanges between the earth’s surface and itsatmosphere. See Figure 9 for an explanation of the units and abbreviations used. If there were a doubling of the CO2 concentration in theatmosphere, one would expect a rise in temperature of 2–3°C at the earth’s surface and in the lower atmosphere. CO2 accounts for about60% of the enhanced greenhouse effect. Note that fossil fuel burning and land-use practices released to the atmosphere about 7,700 Mt ofcarbon in 1995. The atmosphere was a major sink for this anthropogenic carbon, but the ocean and the terrestrial biosphere took up about50% of it. Compare this figure with Figure 4 (after Mackenzie, 1995; Houghton et al., 1996).


26

production and respiration-decay in Figure 11.)This excess organic carbon, approximately 400million tons annually, eventually made its wayinto river waters in particulate and dissolvedforms. It was then transported to the oceans,where some portion of it decayed. The CO2 gen-erated by this decay led to the pristine het-erotrophic state of the ocean. In addition, the pre-cipitation of calcium carbonate in the ocean alsoled to the evasion of CO2 to the atmosphere(Figure 11).

During the past several centuries, as fossilfuel burning and land-use changes added CO2 tothe atmosphere, the ocean went from a net sourceto a net sink of CO2 from the atmosphere. Theconcentration of CO2 in the atmosphere rose, andthe gradient of CO2 concentration between theatmosphere and ocean changed. This changefavored the net uptake of CO2 in the ocean bysolution of the gas in surface seawater (seeEquation 9, p. 27). This oceanic sink of anthro-pogenic CO2 was on the order of 2,000 milliontons of carbon per year during the 1980s,although it varies annually.

Continuing with the exploration of Figure 11,we see that volcanic emissions of CO2 from landand under the sea amount to about 60 milliontons of carbon annually. This flux varies eachyear with the intensity of volcanic activity. Theprocesses of diagenesis and metamorphism cre-ate a flux of about the same size as that of vol-canism. These processes occur as buried sedi-mentary materials change; calcium carbonate andsilica are converted to calcium silicate and CO2.The CO2 is consequently released to the atmo-sphere via volcanism, hot springs, and seepagefrom deep within sedimentary basins.

Weathering of minerals on land removes CO2from the atmosphere, and weathering of sedi-mentary organic matter (kerogen and fossil fuel)at the land surface by oxygen adds CO2 to theatmosphere (Figure 11). As mentioned in Chapter2, the balance between the weathering fluxes andthose of volcanism, diagenesis, and metamor-phism has changed throughout geologic time.These changes are in part responsible for thelong-term variation in atmospheric CO2 duringgeologic time (Figure 6).

Keep in mind, as mentioned in the previoussection, that CO2 is coupled to the reduced car-bon gases through the oxidation of CH4 and COby OH* in the atmosphere. Therefore, any initialwarming of the planet and consequent enhancedemissions of CH4 and CO from the earth’s sur-face because of human activities could potential-ly lead to an increase in the accumulation of CO2in the atmosphere, a positive feedback.

The imbalance in the cycle and feedbacks

The notorious problem with atmosphericCO2 today is the difficulty in balancing theknown sinks with the source from fossil fuelcombustion and land-use activities such as bio-mass burning. The problem is illustrated inFigure 11.

We know that carbon is accumulating in theatmosphere at the rate of 3,300 million tons annu-ally. We also know that the amount of carbon putinto the atmosphere from land-use activities is1,600 million tons of carbon per year (with anuncertainty of 1,000 million tons, plus or minus),and the flux from fossil fuel burning plus cementmanufacturing is 6,100 million tons per year—atotal of around 7,700 million tons. Subtracting theknown atmospheric accumulation of carbon fromthe known source, we see that the fate of roughly3,000 to 4,000 million tons of human-producedcarbon per year is not known. The lower numberis within the rather large error margins of recentestimates of how much anthropogenic CO2 theocean takes up annually, but the upper amount isout of the range of these estimates.

These numbers suggest that there may beanother important sink (sometimes called the“missing sink”) of anthropogenic CO2 besides theatmosphere and ocean. This sink has beenhypothesized to be the Northern Hemisphere’smidlatitude forests and soils. If that is the case, arather strange situation has developed withrespect to the modern carbon cycle. While tropi-cal rain forest ecosystems are sources of carbon tothe atmosphere because of deforestation at ratesof about 1% of the world’s forested area per year,higher-latitude terrestrial ecosystems may besinks of carbon.


27

Such a net sink implies that some processesof carbon storage on land have changed. Some orall of the following changes may have occurred.

• Increased atmospheric levels of CO2 act as afertilizer and stimulate productivity inplants. This leads to storage of carbon in bio-mass or in soil organic carbon.

• Plant productivity is stimulated by increasedNO3

- and NH4+ from fertilizers used in farm-

ing and from deposition of anthropogenic Nfrom the atmosphere. Carbon once more isstored in biomass or in soil.

• Vegetation regrows in previously disturbedecosystems, or grows in undisturbed ones.

The first process on this list, carbon dioxidefertilization, is a potentially strong negative feed-back on the accumulation of anthropogenic CO2in the atmosphere and hence on global warming.We know from experiments on plants and smallecosystems that almost all agricultural crops anda few perennial plants, when subjected toincreased CO2 levels, will increase their rates ofphotosynthesis and growth. If this enhancementshould also occur in large ecosystems, likeforests, CO2 would be withdrawn from theatmosphere and stored in plant organic matter.

Humans have put excess phosphorus, nitrate,and ammonia nutrients into the environment byapplying fertilizers to the land surface, burningfossil fuels and biomass, and discharging sewagecontaining nitrogen and phosphorus. Thesenutrients can stimulate increased plant growth inthe soil and aquatic environments. This eutrophi-cation of both land and marine environments is anegative feedback on accumulation of carbon inthe atmosphere from anthropogenic sources andhence is a negative feedback on warming of theearth. Total land and ocean eutrophication mayamount to a billion tons of carbon per year.

Other processes and feedbacks involving CO2

What about other processes and feedbacks inthe biogeochemical cycle of CO2 that mightchange with a global warming? Perhaps one ofthe strongest potential positive biotic feedbacksin the terrestrial environment is the effect ofincreasing temperature on photosynthesis and

respiration. The rates of both processes in vegeta-tion and microbial life increase with increasingtemperature. However, respiration rates are moresensitive to temperature change. In a warmingworld, increased respiration could temporarilyincrease the flux of CO2 to the atmosphere by asmuch as 1–3 billion tons of carbon per year. Thisincreased flux is a potentially strong positivefeedback on CO2 accumulation in the atmosphereand hence on global warming.

A change in the ocean temperature wouldaffect the amount of dissolved inorganic carbonin seawater. The equation relevant to this effectis

CO2(g) + H2O + CO32- = 2HCO3

- (9)

The (g) is gas, and the equal sign implies thereaction is an equilibrium representation of theprocess.

For a warmer ocean surface, this reactionmoves from right to left—the bicarbonate ionsbreak down into carbon dioxide, water, and car-bonate. Thus the concentration of CO2 in thewater will increase, and CO2 will evade from theseawater to the atmosphere. For a temperatureincrease of 1°C, the change in CO2 concentrationis on the order of 10 x 10-6 atmospheres (10 µatm).This is a positive feedback that could amplifya future atmospheric CO2 increase by about5%.

At a constant temperature, as you add CO2to the ocean, equation 9 goes from left to right,consuming carbonate ion and producing bicar-bonate ion. This process too has the effect ofreducing the ocean’s ability to take up CO2.

Feedbacks involving ocean circulation arestrongly linked to the biogeochemical cycling ofcarbon and nutrients in the ocean. If the oceanbegan to warm, the waters at the surface wouldwarm more than those at depth. This changewould decrease the amount of nutrients risingfrom the deep ocean into the euphotic zone. Withfewer nutrients, there could be a consequentdecrease in biological productivity, followed byless organic carbon escaping the euphotic zoneand settling into the deep sea in a set of processesknown as the biological pump. It is also conceiv-able that changes in wind patterns and windintensities along the western coastal margins of


28

continents could lead to changes in the upwellingof CO2 and nutrients to the surface ocean.

There are several other potential feedbacksinvolving CO2, the ocean, and climatic change,but it is difficult to determine their importance:

• Increased ultraviolet radiation because ofstratospheric ozone depletion may affect thecapacity of certain marine ecosystems, e.g.,around the Antarctic, to remove carbon fromthe atmosphere.

• With warming, the composition and distribu-tion of algae, jellyfish, and other marinespecies could change.

• The rates of delivery to the ocean of iron,molybdenum, and other trace metals essen-tial to marine life may change.

• An increase in the rate of decomposition ofdissolved organic carbon in the ocean couldoccur. There are about 1,000 billion tons ofdissolved organic carbon in the ocean.

• The biological processes associated withthese possibilities could change the concen-tration of CO2 in surface waters and hencethe amount of anthropogenic CO2 the oceancan take up.

• Finally, with a rise in atmospheric CO2 toabout 1,000 ppmv, the pH of the ocean couldfall sufficiently to make it difficult for theorganisms that build skeletons of calciteenriched by magnesium (e.g., coralline algaeand sea urchins) and aragonite (e.g., corals)to produce shells.

29

Nitrogen forms part of the molecules thatmake up living things, such as amino acids (thebuilding blocks of proteins) and DNA. The nitro-gen in proteins bonds together various aminoacids to form the protein structure. The amountof nitrogen in the atmosphere is very large com-pared to that in the oceans or rocks. Of the ele-ments C, N, P, S, and O, only nitrogen is found inmore abundance in the atmosphere than in rocks.

The complete biogeochemical cycle of nitro-

gen is very complex. Figures 12–17 show onlyportions of it. There are six major forms of atmo-spheric nitrogen: the gaseous forms of diatomicnitrogen (N2), ammonia (NH3), nitrous oxide(N2O), and NOx (NO and NO2), and the aerosolsof ammonium (NH4

+) and nitrate (NO3-). In this

chapter, we will focus on the cycles of the firstfour of these forms, and also discuss nonmethanehydrocarbons, the cycles of which are closelyrelated to those of NOx.

The Important Nutrient Nitrogen

Figure 12. Part of the modern global biogeochemical cycle of nitrogen, emphasizing interactions among the land, atmosphere, and ocean.Fluxes between the ocean, land, and groundwater are shown as arrows, with quantities given in Mt N/yr. Fluxes within reservoirs are shownas circling arrows. “Ind. fix” is industrially fixed N (for the manufacture of fertilizers), “Bio. fix” is biologically fixed N, DN is dissolved N, PON isparticulate organic nitrogen, and “pollutant” is the excess nitrogen that has resulted from human activities (modified from Mackenzie, 1995).

NITROGEN(fluxes = Mt N/y)

N2O

N2OAtmospheric

CO2

560

8000

Ocean

Land

Groundwater

42 Rice cultivation20 Combustion78 Ind. fix.126 Bio. fix.

Accumulation

Organic N28

Agriculture9

Human waste20

Aerosol14

River35 DN

27 PON>21 “pollutant”

Enhanced organicproduction-burial

216 Mt C/y

Denitrification1.4 – 2.6

N 2 fix

ation

Evasion



30

N2

The overwhelming majority of nitrogen inthe atmosphere is in the form of N2. The otherforms exist only in small quantities. Biologicalfixation and denitrification are the major process-es leading to exchange of nitrogen between theearth’s surface and the atmosphere (Figure 12).Biological fixation is the process whereby N2 iswithdrawn from the atmosphere and convertedto N compounds that plants can use (e.g., NH3and subsequently NO3

-). Denitrification is theprocess by which nitrogen as N2 or as N2O isreturned to the atmosphere. Both processes aremediated by a variety of bacteria living in soilsand water.

The exchange fluxes between the earth’s sur-face and the atmosphere are small compared tothe internal recycling of nitrogen within the landand ocean realms (Figures 12 and 13). Combustionpractices, the production of commercial fertilizers,

and rice paddy cultivation add fixed nitrogen tothe earth’s surface. Because of these human activi-ties, the amount of nitrogen on land is increasing(Figure 12). About 30 million tons of nitrogen areleached from agricutural fertilizers and humanwaste each year and added to groundwater sys-tems and runoff. Some of this nitrogen makes itsway to rivers and then to lakes and the coastaloceans. On a global scale, rivers may alreadycarry more nitrogen from human activities thanwas transported in the natural state (Figure 12).This increased nitrogen flux to lakes, rivers, andcoastal marine environments is one cause ofincreased regional and global eutrophication ofthese systems. Note, however, that rivers supplyonly a small percentage of nitrogen to the coastalzone (Figure 13). Most of the nitrogen there, otherthan that recycled in the zone, upwelled from thedeep ocean to the surface.

Scientists have calculated how much thishuman-caused increase in nitrogen is likely to be

ORGANIC NITROGEN(fluxes = Mt N/y)

Surfaceopen ocean

Deep ocean

Coastalzone

Verticalmixing

670

Organicmatter

sedimentation900

Export200

Upwelling206

RiversDissolved N 35Particulate N 27

Figure 13. River input of N to the ocean compared to the fluxes involved with the internal recycling of N in the ocean due to biological pro-ductivity and decay. Besides the ocean fluxes shown, about 90% of the nitrogen involved in biological production is simply recycled in theshallow surface waters of the coastal and open oceans. Some nitrogen escapes from the surface ocean in organic matter that settles to thedeep ocean, where the organic matter is decayed and the nitrogen released. It then returns to the surface via upwelling in the coastal zoneand vertical mixing in the open ocean. Some nitrogen, about 30 Mt per year, is buried in marine sediments in organic matter (see Figure12). Some N is transported to the open ocean from the coastal environment (after Mackenzie, 1995; Houghton et al., 1996).

NITROUS OXIDE(fluxes = Mt N/y)

Soils

~7

Com

bust

ion0.

1–

0.3

Biom

ass

burn

ing

0.02

–0.

3

Ferti

lizer

0.01

–2.

2

Ocean

1.5–

3

To s

trat

osph

ere

5 –

9Climate sensitivity0.4°C/N2O doubling

5% greenhouse

N2O1510 Mt N312 ppbv

Residence time = 130 yAccumulation:

3.6 Mt C/y0.8 ppbv

Land Ocean

~ 85%+ hνN2 NOx

<15%+ O3

Figure 14. Part of the modern global biogeochemical cycle of nitrous oxide. Symbols and units are as in Figure 9. This gas is responsible for5% of the enhanced greenhouse effect. A doubling of its atmospheric concentration could lead to about a 0.4°C increase in global tempera-ture. Notice the reaction of this long-lifetime gas in the stratosphere, leading to the destruction of stratospheric ozone. The fluxes in thiscycle are not well known (modified from Mackenzie, 1995).


31

changing ocean productivity and the flux oforganic carbon from the ocean’s euphotic zone.The calculations show that in the 1980s there mayhave been an increased organic carbon flux fromthe atmosphere to the oceanic environment ofabout 200 million tons of carbon per year (Figure12), which is buried in marine sediments. Thisflux takes from the atmosphere about 3% of theincrease occurring today as a result of fossil fuelburning. While relatively small, this is a possiblenegative biotic feedback on atmospheric CO2 andhence global climate change.

N2O

N2O is a natural product of biological deni-trification in soils and in the ocean (Figure 14).The N2O produced by denitrification is onlyabout 15% of all N returned to the atmosphere;the rest is in the form of N2.

N2O is an important greenhouse gas, account-ing for about 9% of the enhanced greenhouseeffect since the 18th century. It has a presentatmospheric concentration of 312 ppmv and a resi-dence time of about 130 years. This concentrationis about 8% greater than in preindustrial time and


32

is increasing at a rate of 0.2–0.3% per year becauseof human activities, including the combustion offossil fuels, burning of biomass, and emissionsfrom urea and ammonium nitrate applied to crop-lands. These emissions amount to 0.13 to 2.8 mil-lion tons of nitrogen annually (Figure 14).

N2O is chemically inert in the troposphere. Inthe stratosphere, it can be converted photochemi-cally to nitric oxide (NO), which acts as a catalystin the destruction of stratospheric ozone (seesidebar). The series of reactions by which this isaccomplished has been one of the regulators ofstratospheric ozone concentration through geo-logic time.

The flux of N2O from the earth to the atmo-sphere has been increasing because of the rapidlyincreasing use of industrially fixed nitrogen (upto the late 1980s), increases in fossil fuel burningand biomass burning, and increases in organiccarbon in coastal waters. This last process is animportant link between the carbon and nitrogencycles. The rate of denitrification and conse-quently of N2O emissions from coastal watersmay have increased because rivers are bringingmore organic carbon to these systems or becausethese systems are undergoing eutrophication asthey receive increased inputs of nutrients fromfertilizer, sewage, and the atmosphere.

With warming, the most important bioticfeedbacks involving N2O are changes in thedenitrification (and nitrification) rates in soils,

sediments, and ocean water. Also, N2O fluxesfrom nitrogen-bearing fertilizers applied to theland surface and sewage discharges into aquaticsystems will be affected by warming. Because thereactions involving N2O are bacterially mediated,it is likely that an increase in temperature willlead to enhanced evasion rates of N2O from theearth’s surface. This is a positive biotic feedbackon accumulation of N2O in the atmosphere and,hence, on global warming. It could also lead to asmall enhanced destruction of stratosphericozone (Figure 14).

NH3

The biogeochemical cycle of ammonia (NH3)is shown in Figure 15. Ammonia is released to theatmosphere by organic decomposition andvolatilization. There, it reacts with water droplets toform ammonium ion (NH4

+) and hydroxyl ion(OH-). NH4

+ appears to be removed from theatmosphere mainly by being deposited back onthe earth in the aerosols of ammonium sulfate[(NH4)2SO4] and ammonium nitrate (NH4NO3).Incidentally, (NH4)2SO4 links the nitrogen and sul-fur biogeochemical cycles, since its deposition onthe earth is also one of the ways oxidized sulfur isremoved from the atmosphere; the other is bydeposition of sulfuric acid (H2SO4).

Two interactions in the NH3 cycle are impor-tant in considerations of global warming. Thefirst is its interaction with OH* to produce NOx.In a warmer world, the decomposition thatreleases NH3 would probably be enhanced,which would slightly increase the stress on theOH* concentration of the atmosphere andenhance production of NOx (Figure 16). Theeffects of increased NOx concentrations are dis-cussed in the following section. The second impor-tant interaction is NH3’s reaction with NO3 andSO4 to produce aerosols containing ammonium(Figure 15). Aerosols are known to cool the planet,although the amount of the effect is unclear. Anincrease in atmospheric NH3 could lead to a smallnegative feedback on potential warming.

The ammonia cycle also gives us informationon nitrogen fertilization of the terrestrial bio-sphere. About four-fifths of the N released to theatmosphere each year in NH3 comes from human

N2O reactions leading to the destruction of stratospheric ozone

NO formation in the middle stratosphere (20–30 km):

N2O + ultraviolet light ⇒⇒ NO + N (10)

N2O + O ⇒⇒ 2NO (11)

Ozone destruction:

NO + O3 ⇒⇒ NO2 + O2

O3 + ultraviolet light ⇒⇒ O2 + O

NO2 + O ⇒⇒ NO + O2

_____________________________________

2O3 + ultraviolet light ⇒⇒ 3O2 (net reaction) (12)


33

activities—50 out of 62 million tons. Only about12 million tons of ammonia nitrogen per yearcomes from natural bacterial decomposition insoils. About 25% of the human-produced flux istransported away from the continents to theoceanic atmosphere. The rest, about 37 milliontons of nitrogen per year, falls back on the landsurface and may be available for terrestrialorganic productivity.

Now it’s time for a back-of-the-envelope cal-culation. If this 37 million tons of nitrogen wereto fertilize land plant production with a ratio of Cto N of 100 to 1, the plants would require morethan 3 billion tons of carbon per year. The phos-phorus accumulating on land each year fromagricultural fertilizers and sewage amounts toabout 8.5 million tons (see Figure 18, p. 36)—justabout the amount of phosphorus needed to

sustain this magnitude of land plant production.This calculation gives some idea of the potentialof fertilization of the land as a sink for the excessCO2 that we are emitting to the atmosphere byfossil fuel burning and land-use practices.

NOx and NMHCs

This brings us to the cycles of NOx and theNMHCs (Figures 16 and 17). We will begin withNOx. It has several natural sources: on the earth,bacterial decomposition of organic matter insoils; in the atmosphere, lightning, mixing fromthe stratosphere, and oxidation of ammonia. NOxalso has anthropogenic sources: fossil fuel andbiomass burning. The main sink of NOx isdeposition on earth of chemical products that wereproduced in the atmosphere by photochemical

AMMONIA(fluxes = Mt N/y) NOx

OceanLand OceanLand

OH* depletion

6

+NH42 Mt N

Variable conc.Residence time =

0.01 y

Orga

nic

deco

mpo

sition

vola

tizat

ion

Com

bust

ion

Organicdecom

position

+ H2O

7

volatilization

33

55W

et-dry

deposition89

NO3

SO4

89

NH32 Mt N

Variable conc.Residence time =

0.01 y

Figure 15. Part of the modern global biogeochemical cycle of ammonia, including that of the ammonium ion (NH4+). See Figure 9 for an

explanation of the units and symbols used. Most of the ammonia emissions from the land surface are due to human activities. Ammonia isremoved from the atmosphere mainly in rain and as small, solid aerosol particles after reaction with water and with nitrate and sulfate.Through reaction with OH*, a small amount of NH3 is converted to nitrogen oxides, e.g., NO and NO2 (modified from Mackenzie, 1995).


34

reactions with NOx, such as HNO3 and organicnitrates.

NMHCs (also called volatile organic com-pounds, or VOCs) are natural byproducts ofplant productivity in terrestrial and marine envi-ronments. Thus, their fluxes to the atmospherechange greatly with the seasons. They also haveanthropogenic sources—once again, fossil fueland biomass burning. Their main sink is in theatmosphere, through oxidation with OH*.

Effects of NOx on ozone

Increasing temperature alone would proba-bly increase the flux of NOx from soils to theatmosphere, potentially depleting OH* and form-ing more methane and ozone in the troposphere(Figure 16). For tropospheric ozone, however, the

effect of increasing temperature is not at allstraightforward. The concentration of ozone inthe troposphere depends in a complex way onthe atmospheric concentrations of several otherbiogenic trace gases, including CH4, CO, and theNMHCs.

In general, when there is little NOx in the tro-posphere (5–30 pptv), increases in the concentra-tions of CH4, CO, and NMHCs lead to a decreasein the concentration of O3. At high NOx concen-trations (generally greater than about 90 pptv),increases in these three gases lead to an increasein ozone. The combination of high NOx andNMHCs in the troposphere disrupts the naturalcycle of production and destruction of ozone,and ozone accumulates. In urban areas, this con-tributes to air pollution.

NITROGEN OXIDES(fluxes = Mt N/y)

NH3

Land OceanLand

Com

bust

ion

21

TroposphericO3

Climate sensitivityGreenhouse gas

OH*depletion6

From

stra

tosp

here

1

Lightn

ing

5

PhotolysisThermal decomposition

Soils

20

Biom

ass

burn

ing

Photochemical

3

Wet-dry

deposition56

56NOx

HNO3PAN

Organic nitrates

Figure 16. Part of the modern biogeochemical cycle of the nitrogen oxides. About one-half of the emissions of these gases to the atmo-sphere comes from the combustion of fossil fuels. In the atmosphere, the gases are converted to several other chemical species, mainlyHNO3 , and removed from the atmosphere both in rain and in dust. Notice the tie to tropospheric O3 (see text), a greenhouse gas and acomponent of smog (modified from Mackenzie, 1995).


35

Effects of NOx on OH*

The concentration of OH*, which is mainlyresponsible for cleansing the atmosphere,depends on the concentrations of trace gases, tro-pospheric ozone, and water vapor. Elevated con-centrations of O3, NOx, and H2O will increaseOH* levels. (Generally, changes in NOx concen-trations affect OH* in the same way they affectozone, described above, except to a lesser degree.)

On the other hand, increases in CH4, CO, andNMHCs will lead to lower levels of OH*.

One critical positive feedback is that increas-es in CO concentrations in the atmosphere couldlead to a reduction in OH*, because NOx has tooshort a lifetime to counteract that effect on a globalscale. Decreased concentrations of OH* could leadto an increase in the lifetime of CH4, a positive,but small, feedback on the accumulation of CH4 inthe atmosphere and hence global warming.

Figure 17. Part of the modern global biogeochemical cycle of the nonmethane hydrocarbons. Land vegetation and phytoplankton naturallyproduce these compounds. Their human sources include industrial practices, transportation, and fossil fuel combustion. These compoundsreact in the atmosphere with OH* and are important in controlling that compound’s concentration in the troposphere. They are also respon-sible for disrupting the natural production and destruction of the ozone cycle in the troposphere. In conjunction with NOx, they can lead toincreased concentrations of O3 in the troposphere (modified from Mackenzie, 1995; Guenther et al., 1995).

NONMETHANEHYDROCARBONS(fluxes = Mt C/y)

TroposphericO3

Climate sensitivityGreenhouse gas

Terre

stria

l veg

eta

tion

terp

ene,

isopr

ene

Biom

ass

burn

ing

150

Ocean

OH*

depletion

1300

+ NOx

1145

com

bust

ion,

solv

ents

ethene, propene

5

NMHCs

Land Ocean

36DRAFT

The Important Nutrient Phosphorus

Phosphorus is a key nutrient, fueling organicproductivity on land and in water. A portion of itscycle is shown in Figure 18. The P cycle is consid-ered here both because it is closely coupled withthose of carbon and nitrogen and for completeness.

Controversy still exists, particularly betweenmarine biologists and geochemists, as to whetherphosphorus or nitrogen is the limiting nutrient inthe marine environment. Globally and on a longtime scale, phosphorus is probably the limiting

nutrient, if for no other reason than that theatmosphere contains essentially an unlimited sup-ply of nitrogen for fixation in aquatic systems.

The major difference between the phosphoruscycle and the carbon, nitrogen, and sulfur cyclesis that no biological process generates an impor-tant gas flux of phosphorus from the earth’s sur-face to the atmosphere (Figure 18). A phosphorusgas known as phosphine or swamp gas (PH3gas)causes a small flux, but the amount is insignifi-cant compared to other fluxes of phosphorus. Aswith nitrogen, fluxes of phosphorus between the

Phosphorus and Sulfur


PHOSPHORUS(fluxes = Mt P/y)

1085

Accumulation

Ocean

186

Land8.5 Acc.

Atmosphere0.03 Mt P

Residence time =2.4 days

Sea salt0.31

Sea salt0.03

Dust3.1

Dust3.72

Dust1.09

Humanactivities

0.47

Wet-dryfallout0.28

Mining12.1

Rivers

Fisheries0.31

2.9 DP (1.5 “pollutant”)

Enhanced organicproduction-burial

140 Mt C/y

Organic P1.6

AtmosphericCO2

Figure 18. Part of the modern global biogeochemical cycle of phosphorus, emphasizing the exchange of P among the land, atmosphere, andocean. Units and symbols are as in Figure 9. Notice, as with nitrogen (Figure 12), that the internal recycling fluxes in the ocean and land reser-voirs by organic production and decay are much larger than the exchanges. Also as with nitrogen, the land is gaining P because of its mining,its use in the manufacture of fertilizers and detergents, and sewage inputs. “Wet-dry fallout” is the precipitation of phosphorus on the ocean asparticles and in rain. DP is dissolved phosphorus, and “pollutant” is excess phosphorus from human activities (modified from Mackenzie, 1995).


37

land, ocean, and atmosphere are small comparedto the amounts that cycle within the land andocean systems.

Sinks and sources

Unlike nitrogen, phosphorus accumulates inboth organic and inorganic sediments. Because ofthe direct tie between N and P in organic matter,the only organic sink shown in Figure 16 is theimportant phosphorus flux to sediments as Pincorporated in dead organic matter (bottomright-hand corner of the figure). The inorganicsinks, which are not shown in the figure, involvethe precipitation of the mineral carbonate fluoro-apatite [Ca5(PO4)3(OH,F)], scavenging by ironoxy-hydroxides, and incorporation in oxidizediron coatings on the surfaces of calcium carbonate.

Phosphorus mining is the largest source of Pto the land surface. Much of the mined phospho-rite rock is used to make commercial fertilizers.The major impact of humankind on the phospho-rus cycle results from the application of both com-mercial and organic fertilizers to croplands andthe disposal of sewage. Phosphorus from thesesources is introduced to streams via direct sewagedischarge or by leaching. The increased amount ofphosphorus in rivers, lakes, and coastal marinewaters results in increased rates of eutrophicationof these systems in cases where phosphorus is thelimiting nutrient.

Feedbacks

Two major biological feedbacks in the phos-phorus cycle may be of importance in global cli-mate change. One involves the enhanced eutrophi-cation of aquatic systems as a result of humanactivities, as just mentioned. The global dissolvedphosphorus flux, carried by rivers to the coasts,has doubled since preindustrial times because ofthese activities (Figure 18). The resulting eutrophi-cation of the coastal regions has led to a potentialaccumulation of organic carbon in these systems ofabout 140 million tons of carbon per year (Figure18, top right). This enhanced flux is a negativefeedback on the accumulation of anthropogenicCO2 in the atmosphere today and, as with nitro-gen, could become more important in the future.

The second biological feedback has to do withthe land reservoir of phosphorus. Notice in Figure18 that phosphorus is accumulating on landbecause of mining, fertilizer use, and sewage dis-charges. Because most chemical weathering andbiological decomposition rates increase withincreasing temperature, a climate warming wouldmean that phosphorus in this reservoir may bemore easily leached into aquatic systems. Thisadditional phosphorus could increase eutrophica-tion in these systems and lead to increased accu-mulation of organic matter. This flux is a negativefeedback on CO2 accumulation in the atmosphereand hence on global warming. However, the effectis quite small. If all the phosphorus now stored onland were leached and transported by rivers tocoastal systems, and all the other nutrients incoastal marine systems could be used for plantgrowth (which may not be the case), the increasein organic carbon storage would be 800 milliontons. This is equivalent to only 13% of one year’sflux of carbon to the atmosphere because of theburning of fossil fuels.

Reduced and Oxidized Sulfur Gases

The reduced and oxidized sulfur cycles(Figures 19 and 20) are closely tied, because thereduced sulfur gases that dominate the earth’sbiological sulfur emissions are oxidized in theatmosphere. These reduced gases are dimethyl-sulfide or DMS [(CH3)2S] and carbonyl sulfide(OCS), which are emitted by the ocean surface,and hydrogen sulfide (H2S), emitted by decayingterrestrial vegetation. Oxidation converts thesegases to sulfur dioxide (SO2) gas and sulfateaerosol, that is, microscopic sulfur-containingparticles. Sulfur enters the atmosphere as theearth emits the reduced sulfur gases and leavesthe atmosphere as sulfate aerosol floats or iswashed to the earth.

The chemistry of DMS and OCS has twomajor connections with climate, by way of cloudcondensation nuclei and stratospheric sulfate.

Cloud condensation nuclei

DMS is produced by bacteria in phytoplank-ton. Its concentration in the oceans is very low, butit is found nearly everywhere at the sea surface,


38

where it may escape into the troposphere. Onceairborne, it is oxidized by OH* to sulfate within afew days. Along with other chemical species in theatmosphere, it condenses into small (micron-sized)aerosol particles. These atmospheric sulfateaerosols act as cloud condensation nuclei (CCN)—the centers on which water droplets may form,facilitating cloud formation. In the remote marineatmosphere, DMS emissions are likely the mainsource of the aerosols that act as CCN.

One hypothesis argues that a warming ofearth’s climate could lead to enhanced phyto-plankton growth and thus greater emission ofDMS from the sea surface. The increased DMSflux could result in increased production of sulfateaerosols and CCN in the remote marine atmo-sphere, creating more and denser clouds. Besides

leading to more rain, clouds also reflect incomingsolar radiation and have a cooling effect on thelower atmosphere and surface of the earth. Thusan increase in cloud cover would give rise to acooling of the troposphere, a negative biotic feed-back on global warming. To counteract a warmingowing to a doubling of atmospheric CO2 wouldrequire a 25% increase in the number of cloud con-densation nuclei.

The validity of this hypothesis is not yet con-firmed by real-world evidence. For example, in icecores collected by drilling at Vostok Station inAntarctica, a record of methane-sulfonic acid(MSA) has been obtained from the ice. MSA is aproduct of DMS oxidation. If this theory were cor-rect, one would expect to find more MSA in icethat dates from warmer periods, since increased

REDUCED SULFUR(fluxes = Mt S/y)

Climate sensitivityDMS→SO4→CCN→cooling

OH* depletion67

Sulfate veil

Biol

ogic

alde

cay

Decay-excretion

SO4

25DM

S,H 2S

42DM

S, OCSTo

str

atos

pher

e

0.06

OCS

SO2

Reduced S0.5 Mt S0.1 ppbv


Land Ocean

Figure 19. Part of the global biogeochemical cycle of reduced sulfur. See Figure 9 for an explanation of the units and symbols used. Reducedsulfur gases are naturally produced both on land and in the ocean by biological processes. They are eventually removed from the atmo-sphere by reaction with OH* and conversion to SO2 and thence to sulfate aerosol (see Figure 20). Sulfate aerosol is produced by the oxida-tion of OCS in the stratosphere to form the sulfate veil during times of low volcanic activity. Sulfate aerosol is responsible for the scatteringand reflection of solar energy and thus cools the planet (modified from Mackenzie, 1995).


39

levels of the compound would wash out of theatmosphere. However, the record shows just theopposite: higher levels of MSA in the ice from thelast ice age, which culminated 18,000 years ago,than in this and past interglacial stages.

Stratospheric sulfate

OCS is produced mainly by the photolysis oforganic sulfur in the surface waters of the oceanand by photochemical oxidation of the biogenicgas carbon disulfide (CS2) in the atmosphere. It isthe most abundant reduced sulfur species in theremote marine atmosphere. Because it is chemi-cally inert, it has a long residence time in the tro-

posphere and can enter the stratosphere. Once inthe stratosphere, OCS is destroyed by reactingwith ultraviolet light and atomic oxygen. It is con-verted to SO2 and then on to sulfate aerosol. Thisreduced sulfur gas supplies about half of themass of sulfate aerosol found in the lower strato-sphere (the so-called Junge layer).

If global warming causes any change in theflux of OCS to the stratosphere, it will have aneffect on climate through change in the Jungelayer. An increased stratospheric sulfate burdenwould give rise to cooling of the atmosphere; adecreased burden would lead to warming. Thelikely feedback effect on an initial warming is dif-ficult to predict.

Figure 20. Part of the global biogeochemical cycle of oxidized sulfur. See Figure 9 for an explanation of the units and symbols used. Noticethat the flux of oxidized sulfur from land as SO2 is dominated by the combustion of fossil fuels and biomass burning. This sulfur reacts withOH* to produce sulfate aerosols of ammonium and hydrogen. About 40% of the sulfur falling back on the earth’s surface is derived fromthese sources. Volcanism can add sulfate aerosols to the stratosphere and produce a temporary cooling of the planet, as after the eruptionof Mt. Pinatubo in 1991. Sulfate aerosols derived from combustion of fossil fuels and biomass burning exert a strong cooling effect, but onlyin the regions where they are produced (modified from Mackenzie, 1995).

OXIDIZED SULFUR(fluxes = Mt S/y)

Reduced S

Ocean

Biom

ass

burn

ing

7 Rain and drydeposition

OH* depletion

67

Climate sensitivityCooling

Volc

anis

mTo

str

atos

pher

e

Sulfate veil

To troposphere

OH* depletion

49

Seaaerosol

115

Com

bust

ion

80

Volc

anis

m10

160

45

SO4

SO20.5 Mt S0.1 ppbv


Land Earth surface

SO4 aerosol2 Mt S

0.35 ppbvResidence time =

4.6 days


40

Oxidized sulfur

To complete the picture of the global cycle ofsulfur gas species, the earth surface–atmosphereoxidized sulfur cycle is shown in Figure 20.Natural sources of oxidized sulfur in the atmo-sphere include oxidation of reduced sulfur gasesas mentioned above, volcanism, and aerosols fromthe sea. The oxidized sulfur is removed from theatmosphere by deposition of aerosols.

One of the most dramatic demonstrations ofthe connection between volcanism, sulfateaerosols, and climate was the eruption of Mt.Pinatubo in the Philippines in 1991. The eruptionplume rose high into the stratosphere, where sul-fate aerosol was generated and distributed overmuch of the globe. The aerosol scattered andreflected solar energy back to space. This event ledto a cooling of the planet of about 0.5°C during1991–93. At its maximum in 1993, the coolingreached –3–4 watts per square meter. This is con-siderably more than the enhanced greenhouseforcing of +2.5 watts per square meter.

Human activities have strongly interferedwith the global biogeochemical cycle of oxidizedsulfur. This disturbance in the cycle has led to theacidification of land and freshwater systems—the

problem of acid precipitation. Of the total globaldeposition of oxidized sulfur today, approximately40% is derived from the human activities of com-bustion of fossil fuels, biomass burning, andsmelting of sulfide ores. The deposits from theseactivities vary enormously spatially.

Recently a new hypothesis has emerged link-ing our sulfur emissions with climate. In theNorthern Hemisphere, emissions of SO2 have ledto an increased mass of sulfate aerosol. Because ofthe aerosol’s short lifetime in the atmosphere, itremains concentrated over the eastern half of thecontinent and the western Atlantic Ocean. Thisregional enhancement could be cooling the atmo-sphere enough to explain the discrepancy betweenthe observed temperature record of the past 100years and the higher temperatures predicted byclimate models that include only the effects ofincreased greenhouse gases and not increasedaerosols. On the other hand, when climate model-ers attempt to introduce the cooling effect ofsulfate aerosols into their models, the modeledtemperatures are generally lower than thoseobserved in the recent past. This difficulty impliesthat sulfate aerosols are an important componentof the temperature record, but that the amount bywhich they cool the planet is not yet known.

41

The water cycle (Figure 21) is important in thecontext of biogeochemical cycling in the C, N, P, S,and O system, as well as being important in itsown right. Water circulating through the ecosphereis part of a continuous hydrologic cycle that makeslife on earth possible. The water cycle is the drivercycle for transport of many elements at the earth’ssurface. In the atmosphere, water vapor is the mostimportant greenhouse gas, and its behavior duringa global warming is of concern.

The water cycle is a balanced system, withwater stored in many places at any one time(Table 3). The cycle involves the transfer of waterin its various forms of liquid, vapor, ice, andsnow through the land, air, and water environ-ments. Both matter and energy are involved inthe transfer. The transfer begins when heat fromthe sun warms the ocean and land surfaces andcauses water to evaporate. The water vapor entersthe atmosphere and generally moves with the

The Water Cycle

Atmospheric water vapor transport(40)

Solar radiation

Cloud

Evaporation

Transpiration

(71)

Lake

Surface water

Soils

Evaporation(425) Precipitation

(385)

Ocean

Groundwater flow

Land

Precipitation ofSnow, ice, rain

(111)

Surface runoff

Percolation

Glacier

River

Return flow (40)

Figure 21. The global biogeochemical cycle of water (the hydrologic cycle), showing the major processes of water movement.Numbers in parentheses show the water budget in thousands of cubic kilometers per year (after Maurits la Rivière, 1990.)



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circulation of the air. On a globalscale, warm air rises in the atmos-phere and cooler air descends. Thewater vapor rises with the warmair. The farther from the warm plan-etary surface the air travels, thecooler it becomes. Cooling causeswater vapor to condense on smallparticles (cloud condensationnuclei) in the atmosphere and toprecipitate as rain, snow, or ice andfall back to the earth’s surface.When the precipitation reaches theland surface, it is evaporated direct-ly back into the atmosphere, runsoff or is absorbed into the ground,or is frozen in snow or ice. Also,plants require water and absorb it,retaining some of the water in theirtissue. The rest is returned to the atmospherethrough transpiration. Precipitation on land isequal to evapotranspiration plus runoff to theocean. That is, over the land, there is more pre-cipitation than evapotranspiration.

For the ocean, the situation is reversed. Muchof the water evaporated from the ocean returnsthere directly; however, a small amount (about8% of that evaporated) is carried by atmosphericwinds over the continents, where it precipitates.Once on the ground, the water finds its way tostreams, lakes, or rivers in runoff or by percola-tion into and through groundwater. In due time,the water will return to the ocean, mainly instream and river flows and less importantly ingroundwater. This return flow balances the netloss from the ocean surface by evaporation.

Snow and ice may remain on the land for along time before the water in these forms of pre-cipitation evaporates to the atmosphere orreturns via rivers or as direct glacial meltwater tothe oceans. The snow and ice that feed glaciersmay remain locked up in the cryosphere forthousands of years, but finally the ice will melt,and the water will travel to another part of thehydrologic cycle.

Water vapor in the atmosphere is the princi-pal greenhouse gas. Because the amount of watervapor in the atmosphere is dependent on thetemperature of the planet, any initial warming of

the planet will probably lead to more watervapor in the atmosphere. The increased watervapor has the potential to absorb more infraredradiation reradiated from the planetary surfaceand thus lead to further warming. This is anotherexample of a positive feedback.

Water droplets form in the troposphere bycondensation of water on cloud condensationnuclei. These particles may absorb or reflect ener-gy. The amount of water vapor in part deter-mines the types and distribution of clouds thatform in the atmosphere. In terms of predictingthe effects of increasing concentrations of green-house gases in the atmosphere on temperatureand other climatic variables, general circulationmodels (GCMs) and other types of models havebeen used. The GCMs are very complex comput-er representations of the atmosphere or atmo-sphere-ocean system that are used in the model-ing of global climate change. In these models, theeffects of clouds on the radiant energy budget ofthe planet are a major source of uncertainty inattempting to predict future climate change.

Clouds regulate the radiative heating of theplanet. They reflect a significant part of theincoming solar radiation. Clouds also absorblongwave, infrared radiation emitted by the earth.At the cold tops of clouds, energy is emitted tospace. In 1984 the Earth Radiation BudgetExperiment (ERBE) was launched. This

Table 3. Distribution of water in the ecosphere

Reservoir Volume (106 km3) Percent of total________________________________________________________________Ocean 1370 97.25

Cryosphere(ice caps and glaciers) 29 2.05

Groundwater 9.5 0.68

Lakes 0.125 0.01

Soils 0.065 0.005

Atmosphere 0.013 0.001

Rivers 0.0017 0.0001

Biosphere 0.0006 0.00004___________________________________________________________

Total 1408.7 100

After Berner and Berner, 1996.


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experiment involves a system of three satellitesthat provide data on incoming and outgoing radi-ation. One result of this experiment so far is thedemonstration that clouds have a net coolingeffect on the earth. On a global scale, cloudsreduce the amount of radiative heating of theplanet by –13 watts per square meter. This is alarge number when compared to the +2.5 wattsper square meter attributed to the increase inatmospheric greenhouse gases during the lastcentury. It is also large compared to the radiativeheating that could arise from a doubling ofatmospheric carbon dioxide concentrations in thenext century—about +4 watts per square meter.Thus, a small change in the types and distributionof clouds may have a large effect on the radiationbudget compared to the effect of changing green-house gas composition owing to human activities.

In a world already warmed by greenhousegases released from human activities, it is difficultto predict what will happen to the types and dis-tribution of clouds. Too little is known about thecomplex processes of cloud formation and, in par-ticular, the response of clouds to a warming earth.Also, these complex processes are difficult to sim-ulate in the general circulation models. One diffi-culty lies with the problem of reproducing cloudphysics in the models. A second difficulty is that

GCMs only perform calculations at widelyseparated points over the globe and relativelyinfrequently. Cloud formation involves verydynamic processes at short time and spatialscales.

Clouds may act as a positive or negative feed-back in a future earth warmed by the enhancedgreenhouse effect. This ambiguity accounts inpart for the range of estimates of the averagetemperature increase predicted by the GCMs fora doubling of the atmospheric carbon dioxideconcentration.

A final comment on the water cycle is that itis being significantly affected by water usage andcontamination of water stocks. Currently, humansuse an amount of water equivalent to about 25%of total terrestrial evapotranspiration and 55%, or6,800 cubic kilometers per year, of the runoff ofwater from the continents that is accessible. Onlyabout 20% of the world’s drainage basins havepristine water quality. There is little doubt thatthe world’s water resources will be significantlystrained in the 21st century.

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Conclusion


An analogy is commonly made between theearth’s surface system and a giant chemical engi-neering factory. In the natural system, materialcirculation is driven by energy from the sun and,to a much lesser extent, from radioactive decay ofelements in the earth’s interior and motions of itstides. This is a mechanical and inorganic view ofthe earth. In another and more realistic sense,the earth has a natural metabolism; materialshave circulated about its surface for millions ofyears in a complex, interconnected web of bio-geochemical cycles. An array of physical, chemi-cal, and biological processes weather and eroderocks and transfer materials in and out of theatmosphere, from the atmosphere to the biotaand back again, to the oceans via rivers, and tothe continents by uplift. Each element has a nat-ural biogeochemical cycle. It is these cycles andtheir relationship to the physical climate systemthat have led to the development of a relativelystable and resilient surface system during

geologic time. Life has evolved in this systemand plays a strong role in the development andmaintenance of the system through processes,fluxes, and feedbacks.

Human activities have contributed materialsto the biogeochemical cycles. Some of these mate-rials enter element cycles already naturally inoperation; they are the same chemical species thathave circulated for millions of years. Other mate-rials are synthetic compounds and are foreign tothe natural environment. These anthropogenicfluxes are leading to a number and variety ofenvironmental issues, including the possibility ofglobal climate change. There is no doubt thathuman activities have interfered in biogeochemi-cal cycles and have modified the composition ofthe atmosphere. Understanding the consequencesof this interference requires better quantitativedescriptions of these cycles, their interconnec-tions, and, in particular, their coupled response toperturbations, such as a change in climate.

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Questions

Chemistry

1. The most simple chemical expression for the production of organic matter in plants is

CO2 + H2O = CH2O + O2

The chemical compound CH2O is organic matter; CO2 is carbon dioxide gas; O2 is diatomic oxygen; and H2O is water. The atomic weights of the elements C, H, and O are, respectively, 12, 1, and 16.

A. What are the gram molecular weights of the compounds of CH2O, CO2, O2, and H2O?

B. What is the weight of one mole of each of these compounds?

C. If 10 moles of plant matter have been produced, how many moles of CO2 did it take to producethe plant matter? How many grams of CO2?

D. If the C:N ratio of the plant material in moles is 106:16 (that of marine phytoplankton), how manymoles of nitrogen were needed to produce the plant matter?

E. If the source of the nitrogen were nitrate (NO3-) in the euphotic zone of the ocean, how many

grams of nitrate were consumed in the process?

2. Write a balanced chemical reaction for the weathering of the mineral orthoclase feldspar (KAlSi3O8)by water containing dissolved CO2. The products that are formed in this weathering reaction are:kaolinite [(Al2Si2O10(OH)4], bicarbonate (HCO3

-), monomeric silicic acid (H4SiO4o), and dissolved

potassium ion (K+). (Hint: to begin, balance the reaction on aluminum.)

3. The concentration of dissolved potassium (K+) in the ocean is 390 mg/kg. The atomic weight of potassium is 39. The average density of seawater is 1.027 g/cm3.

A. What is the concentration of K+ in seawater in moles/kg?

B. What is its concentration in parts per million by volume of seawater?

4. The average molecular weight of the gases in the atmosphere is 29 and the mass of the atmosphere is 52 x 1020 grams.

A. What is the total number of moles of gases in the atmosphere?

B. Carbon dioxide gas makes up 0.036 % of the atmosphere in moles. How many moles of CO2are there in the atmosphere? How many grams?

Study Questions and Answers



46

5. Water vapor in the atmosphere averages about 0.2% of the atmosphere in weight.

A. What is the total mass of water vapor in the atmosphere?

B. How many moles of H2O vapor are there in the atmosphere?

The Oceans, Atmosphere, Sediments, and Rocks

1. The average depth of the ocean is 3.8 kilometers and the average upwelling rate of deep water into thesurface open ocean and into coastal environments is 4 m/yr. The area of the ocean is 3.6 x 1018 cm2.

A. About how long would it take to upwell the entire volume of the ocean?

B. The average nitrogen content of deep water is about 40 x 10-6 moles per liter (L). What is theannual rate of addition of nitrogen to the surface water due to upwelling?

C. At a molar ratio of C:N of 106:16, what is the productivity in the global surface ocean in grams (g)C/m2/yr sustained by the upwelling of nitrogen?

2. One source of the deep water of the world’s oceans is in the Norwegian Sea. Here water is sufficient-ly dense to sink to the bottom. This water mainly forms from the cooling by evaporation of watercarried northward by the Gulf Stream. The Gulf Stream carries heat to the high latitudes of theNorth Atlantic Ocean which helps to moderate the climate of Europe. There is considerable interestin the rate of formation of North Atlantic deep water (NADW) because of the role it plays in climate.Possible changes in the rate of its formation have been cited as the cause of rapid climate change inthe past. Any global warming could modify the rate of deep water formation.

A. How would you expect the residence time of the deep water to change from the Atlantic Ocean tothe Pacific Ocean?

B. How does the deep water return to the surface?

C. If the continental glaciers were to begin melting because of a global warming, what would youexpect might happen to the rate of deep water formation and the climate of Europe?

3. The winds of the earth tend to blow along latitudinal lines around the planet. The major wind belts arethe Polar Easterlies, Westerlies, Trade Winds, and Equatorial Easterlies in both the Northern andSouthern Hemispheres. The winds drive the surface currents of the ocean. Cool air that has descendedat midlatitudes moves toward the equator as the Northeast and Southeast Trade Winds in the Northernand Southern Hemispheres, respectively. These winds exert a force on the sea surface, and currents aregenerated that flow toward the equator. Both the winds of the atmosphere and the surface currents ofthe ocean converge near the equator. The warm equatorial air rises and moves north and south towardthe poles. The converging ocean currents generate westward-flowing equatorial currents.

A. Based on the above, would you expect a pollutant gas like carbon monoxide with major anthro-pogenic sources in the Northern Hemisphere and a short atmospheric residence time to be evenlydistributed in the troposphere?

B. Is there an equatorial barrier to the dispersal of floating tar balls produced when petroleumtankers spill oil in the shipping lanes of the North Atlantic?

C. Does such a barrier exist for the deep waters of the ocean?


47

4. The very finest particles of airborne dust carried by winds off the Sahara Desert travel in the tropo-sphere for long distances westward across the Atlantic Ocean. These particles are deposited on theocean surface and settle out at a rate of 500 cm/yr. How long would it take such particles to reachthe bottom of the Atlantic Ocean at 4 km?

5. The global atmosphere in 1996 contained about 360 ppmv of CO2. This concentration is equivalent toa partial pressure of CO2 (PCO2) of 10-3.45 atmosphere (atm).

A. What is the concentration of dissolved CO2 in the surface ocean in equilibrium with atmosphericCO2 at a temperature of 25°C?

B. In the year 1700, when atmospheric CO2 was at a concentration level of 280 ppmv (PCO2 = 10-3.55

atm), what was the concentration of CO2 in the surface ocean?

C. What was the percentage change in atmospheric CO2 concentration between 1700 and 1996?

6. The primary energy sources for the earth are:

• solar radiation, 0.5 cal/cm2/min (about 343 W/m2)

• heat flow from the interior of the earth, 0.9 x 10-4 cal/cm2/min

• tidal energy, 0.9 x 10-5 cal/cm2/min.

About 49% of the solar radiation is absorbed by earth’s surface and reradiated to space as longwave,infrared radiation.

A. What percentage of the total comes from the combination of heat from the interior of the earthand tidal energy?

B. In units of W/m2, how much solar radiation reaches the earth’s surface and is absorbed there?What does this energy do?

C. Of the 388 W/m2 of longwave radiation emitted to space by the earth’s surface, 326 W/m2 areabsorbed by water vapor, CO2, and other greenhouse gases in the atmosphere and reradiatedback to the earth. What happens to this energy, and what effect does it have on the earth?

7. The area of land today is 150 x 1012 m2, and the mean elevation of the continents is 0.84 km. The con-tinents are being eroded at a rate of 200 x 1014 g/yr. The average density of rock is 2.7 g/cm3.

A. At the current rate of erosion, how long would it take to wear the continents down to sea level?

B. Your answer to question 7A was a period of time much less than the age of the earth of 4.6 billionyears. In that case, why are there any continents?

8. The total mass of sedimentary rock younger than 600 x 106 years is 1.8 x 1018 metric tons. Its meanresidence time is 400 x 106 years.

A. In an unchanging system, what is the flux of sediment in and out of the sedimentary rock reservoir?

B. How does sediment get into the sedimentary rock reservoir and how does it get out?


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9. What chemical species of nitrogen, sulfur, and carbon might you expect to find in a water-logged(anoxic) soil? In a well-aerated soil?

10. What does the weathering of silicate minerals subtract from or add to the atmosphere?

Ecology

1. Tropical rain forests have a total area of 17 x 1012 m2, and estuaries have an area of 1.4 x 1012 m2.Their mean net primary production is 2,000 g dry plant matter/m2/yr and 1,800 g dry plantmatter/m2/yr, respectively. Their mean plant biomasses in kg C/m2 are 20 and 0.45, respectively.Forty-five percent of dry plant matter is carbon.

A. What is the total net primary production of tropical rain forests and estuaries in metric tons of dryplant matter and carbon per year?

B. What is the total plant mass of these ecosystems in metric tons of dry plant matter and carbon?

C. The tropical rain forests of the world lost 9% of their area due to cutting in the 1980s. How muchdry plant matter does this cutting represent? How much carbon?

D. If all of the carbon in the cut trees of question 1C were emitted to the atmosphere by burning andslow oxidation, how many grams of CO2 would this represent? What fraction of the atmosphericCO2 reservoir is this (see Figure 11)?

E. Estuaries receive 1.5 x 1012 g of pollutant dissolved phosphorus annually (see Figure 18). Thisamount of P could support how much additional plant productivity as dry matter and as C/m2

per year?

F. If all the additional plant matter in 1A were buried in the sediments of the estuaries, would thisflux qualify as a biological feedback to the accumulation of CO2 in the atmosphere? Explain. Whatpercentage of the atmospheric flux of 6 x 1015 g C from fossil fuel burning in 1995 does the addi-tional total plant production in estuaries represent?

2. The net primary production of the earth’s surface is 0.37 kg dry matter/m2/yr.

A. With a total area of 510 x 1012 m2, what is the total production of dry matter?

B. How much carbon does the production in 2A represent?

C. Total production on land exceeds that in the ocean by perhaps as much as two times. What is theannual total production on land and in the ocean?

3. What is the difference between autotrophy and heterotrophy?

4. What are three biogeochemical processes performed by the prokaryotes?

5. What is cultural eutrophication? Why may this phenomenon qualify as a negative feedback to theaccumulation of CO2 in the atmosphere?


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Biogeochemical Cycles

1. Construct a diagram showing the reservoirs and fluxes in the global biogeochemical cycle of water.Use the information in Figure 21 and Table 3, and the reservoirs of atmosphere, land, and ocean.

A. What is the residence time of water in the ocean with respect to the flux of water via the rivers tothe ocean?

B. If the concentration of dissolved calcium is 400 ppm in the ocean and 15 ppm in average riverwater, what is the residence time of calcium in the ocean?

2. A nearly rectangular-shaped lake is 5 km long, 2 km wide, and 100 m deep and contains 0.001 mg/Lof dissolved mercury. A river discharges 2 x 1012 L/yr of water with a concentration of mercury of0.0005 mg/L into the lake.

A. What is the water volume in the lake (in liters)?

B. What is the total mass of mercury in the lake?

C. What is the residence time of mercury in the lake?

D. An industrial plant near the mouth of the river accidentally discharges mercury into the lake. Howlong will it take for most (95%) of the mercury contamination to work its way out of the lake?

3. The terrestrial living biomass contains 600 x 109 tons of carbon. It is argued that during the decade ofthe 1980s, about 135 x 1012 moles of anthropogenic CO2 were taken up annually by the terrestrialbiomass.

A. What would be the average annual percentage increase in the mass of carbon in the terrestrialbiosphere?

B. Do you believe such an increase would be detectable by doing field studies to determine theincrease in biomass?

4. The mixed layer of the ocean is the vertical layer that is well stirred by winds blowing across the sur-face of the sea. Chemical and physical characteristics of the water column are rather uniform overthe depth of the mixed layer. The average thickness of the mixed layer throughout the world’soceans is about 100 meters but varies from 50 to 300 meters. The average total dissolved inorganiccarbon (DIC, HCO3

- + CO32- + CO2) content of the mixed layer is 2.2 micromoles (mmol) per kg.

A. What is the total mass (reservoir) of DIC in the mixed layer of the ocean in moles of carbon? (Thearea of the ocean is 360 x 1012 m2.)

B. The ocean takes up about 2 billion tons of carbon annually from the human activities of fossil fuelcombustion, cement manufacturing, and deforestation. What is the annual increase in dissolvedcarbon in the mixed layer of the ocean in mmol/kg due to this absorption of anthropogenic CO2?

C. Use the annual increase you derived from 4B and assume the DIC content of the mixed layer was2.2 mmol/kg in 1975. What would be the percentage increase in the DIC content of the mixedlayer from 1975 to 1996?


50

5. Total evaporation is 496 x 103 km3 H2O/yr over the earth’s surface.

A. What is the residence time of water in the atmosphere?

B. If a pollutant is susceptible to being rained out (washed out) of the atmosphere, would you expectit to mix evenly throughout the troposphere?

6. What are the three major processes controlling atmospheric CO2 concentration levels on a long timescale? Write a balanced set of chemical equations demonstrating these processes.

7. What is the connection between organic matter and atmospheric O2?

8. What is the major difference between the biogeochemical cycle of phosphorus and those of carbon,nitrogen, and sulfur?

9. What is the major difference in the reactivity of DMS and OCS in the atmosphere?

10. The mean atmospheric lifetime (residence time) of NH3 is 14 days; that of CO is 60 days. How do therates of destruction of these gases in the atmosphere by OH* compare qualitatively?

11. Referring to Figures 9–11, what is the major reaction in the atmosphere that couples the biogeochemicalcycles of CH4, CO, and CO2? What percentage of the enhanced greenhouse forcing on climate is due toCO2? To CH4? What are the sources of emissions of CH4 to the atmosphere from human activities? OfCO2 and CO? If the concentration of CO2 in the atmosphere were doubled, what would be the poten-tial increase in temperature? How much more effective is CH4 than CO2 as a greenhouse gas?

12. Referring to Figure 12, what is the ratio of anthropogenic N fluxes on land involving fixation ofatmospheric N to natural biological fixation? What is the minimum percentage of the N fixed byhuman activities that is discharged to the ocean by rivers? What is the important environmentalproblem related to this additional N flux to the ocean?

13. Referring to Figure 13, what is the ratio of these two fluxes: the upwelling flux of N to the coastal zone andthe dissolved N flux to the ocean via rivers? Which flux would you expect to change during the nextcentury?

14. Referring to Figure 14, based on the summation of the high and low estimates of fluxes of N2O fromearth’s surface to the atmosphere, what is the range in residence time estimates for N2O in theatmosphere? What percentage of the enhanced greenhouse forcing of climate is due to N2O?

15. Referring to Figure 16, of the total flux of NOx to the atmosphere, what percentage is from humanactivities? What is the principal type of reaction that destroys NOx in the atmosphere? Write thereaction for the destruction of NO2 and its subsequent removal as HNO3.

16. Referring to Figure 18, the total mass of P in land plants is 1,800 million tons and that in marineplankton is 73 million tons. What is the ratio of the internal recycling flux in the ocean to that on land?What is the residence time of P in the land biota relative to the recycling flux? In the marine plankton?


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Answers

Chemistry

1. This question is designed to enable the student to review basic knowledge of the meaning of a chem-ical equation. Many biogeochemical processes are simply described by chemical reactions. In thisreaction, there are four pure substances that can be decomposed by a chemical change, that is, fourchemical compounds: CO2(gas), H2O(liquid), CH2O(solid), and O2(gas). The atoms in these compoundsmay separate from one another and rearrange themselves during a chemical reaction. In this case, inthe presence of light and available nutrients, plant material has formed from CO2 and H2O, and thecarbon atom in CO2 has been rearranged into the carbon atom of CH2O. This reaction further showsthat all chemical equations must balance. The total number of atoms of an element on the left-handside of an equation must equal the total number of atoms of that element on the right-hand side ofthe equation.

A. The atomic weights of the elements in these compounds are C = 12, H = 1, O = 16. Thus, the sum-mation of these weights, or the gram molecular weight, for CH2O is 12 + (2 x 1) + 16 = 30 g; forO2, 2 x 16 = 32 g; for CO2, 12 + (2 x 16) = 44 g; and for H2O, (2 x 1) + 16 = 18 g.

B. The weight of one mole of these compounds is CH2O = 30 g; O2 = 32 g; CO2 = 44 g; and H2O = 18 g.

C. The relationships between the compounds as shown in the equation must always be preserved;thus one unit—that is, one mole of the reactant compound CO2—must react with an equivalentnumber of moles of the reactant compound H2O to give you a ratio of product compounds ofCH2O:O2 = 1:1. Thus, the formation of 10 moles of plant material requires 10 moles, or 10 moles x44 g/mole = 440 g of CO2.

D. At a molar ratio (that is, a ratio of moles) of C:N = 106:16, the moles of nitrogen are 16/106 x 10moles = 1.51 moles.

E. The atomic weight of N is 14. The reaction required 1.51 moles of N or 1.51 moles of NO3-. (There

is 1 mole of N in 1 mole of NO3-.) The molecular weight of NO3

- is 14 + (3 x 16) = 62 g. Thus, thegrams of nitrate consumed were 1.51 moles NO3

- x 62 g NO3-/mole = 93.6 g NO3

-.

2. This question again deals with chemical reactions. The equation to be balanced represents the weath-ering of a common mineral found at the surface of the earth. The student not only learns to balance areaction on the basis of atoms but also on the basis of charge. The reason for starting the balance onaluminum is that solids containing aluminum are very insoluble at the temperature and pressure ofweathering. Thus, it is assumed that all the aluminum from the weathering of orthoclase feldspar issimply transferred to the solid weathering product kaolinite. Of course, this is not true, but it is areasonable approximation.

2KAlSi3O8 + 2CO2 + 11H2O = Al2Si2O5(OH)4 + 2 K+ + 2HCO3- + 4H4SiO4

0

3. Dissolved potassium is the sixth most important dissolved constituent in seawater. Its concentration isimportant to biological processes, although it is a minor essential element for phytoplankton productivity.

A. The concentration of K+ in the ocean in moles/kg = 390 x 10-3 g/kg seawater ÷ 39 g/mole = 10 x10-3 moles/kg seawater or 10 millimoles.


52

B. The average density of seawater is 1.027 g/cm3. Thus, 1 kg of seawater is equivalent to 1,000 g ofseawater ÷ 1.027 g/cm3 = 973.7 cm3 of seawater. The concentration in parts per million by volumeof seawater is 390 x 10-3 g/kg seawater x 1.027 x 10-3 kg/cm3 x 1,000 cm3/L seawater = 400.5 x10-3 g/L = 400.5 mg/L = 400.5 ppmv of seawater.

4. The atmosphere is an important reservoir that exchanges materials with the earth’s surface. Its composi-tion has changed on a geological time scale and on the human scale of generations because of the activi-ties of human society. This problem illustrates the fact that although atmospheric CO2 accounts for 60% ofthe enhanced greenhouse effect, it represents a relatively small part of the total atmospheric mass.

A. The mass of the atmosphere can be obtained from the weight of air above 1 cm2 of the earth’s sur-face and the total area of the earth: 1,031 g/cm2 x 5.1 x 1018 cm2 = 52 x 1020 g. The total number ofmoles of gases in the atmosphere is 52 x 1020 g ÷ 29 g/mole = 1.8 x 1020 moles.

B. The number of moles of CO2 is 0.00036 x 1.8 x 1020 moles = 64.8 x 1015 moles. The mass of CO2 ingrams is 64.8 x 1015 moles x 44 g/mole = 2,850 x 1015 g = 2,850 x 109 metric tons = 2,850 gigatonsx (12 ÷ 44) = 777 gigatons of C ÷ 12 g/mole = 64.8 x 1015 moles C.

5. Water vapor is the most important greenhouse gas, yet as the following calculation shows, it is asmall portion of the mass of the atmosphere.

A. The total mass of water vapor in the atmosphere is 0.002 x 52 x 1020 g = 10.4 x 1018 g.

B. There are 10.4 x 1018 g ÷ 18 g/mole = 57.8 x 1016 moles of water vapor.

The Oceans, Atmosphere, Sediments, and Rocks

1. Upwelling is an important process in the ocean. It involves the upward movement of water and dis-solved constituents from depth in the ocean to the surface. Upwelling occurs in the coastal regions ofoffshore Peru, California, Namibia, Mauritania, and Somalia, and in open ocean equatorial regionsand the high latitudes of the Southern Hemisphere.

A. To upwell the volume of the ocean with an average depth of 3,800 m at the mean upwelling rateof 4 m/yr would take 3,800 m ÷ 4 m/yr = 950 years.

B. 400 cm of water rises 1 cm2/yr. With an ocean area of 3.6 x 1018 cm2, this is equivalent to a vol-ume of water of 1,440 x 1018 cm3/yr or 1,440 x 1015 L/yr x 40 x 10-6 moles N/L = 57.6 x 1012

moles N/yr x 14 g/mole = 806 x 1012 g N/yr.

C. 806 x 1012 g N/yr ÷ 14 g/mole = 57.6 x 1012 moles of N/yr. At a C:N ratio of 106:16 = 382 x 1012

moles C/yr x 12 g/mole = 4.6 x 1015 g C/yr ÷ 360 x 1012 m2 = 13 g C/m2/yr. This productivity isonly about 10% of global marine net primary productivity (Figure 11). This calculation illustratesthe fact that much of the nitrogen used in biological productivity in the euphotic zone of theocean comes from recycling of the N within this zone.

2. The formation of the deep water of the ocean is part of the conveyor belt circulation pattern of theworld oceans. This pattern is not well known from observations; our understanding of it is basedmainly on theoretical models. The North Atlantic deep water (NADW) flows southward at depthfrom its source in the high latitudes of the North Atlantic Ocean and meets a northward-flowing cur-rent (the Antarctic Bottom Water, ABW) whose water originated by sinking in the Weddell Sea nearAntarctica. The currents merge and part of the water flows into the deep Indian Ocean and Pacific


53

Ocean. Water upwells to the surface in all basins of the ocean. In addition, warm water returns fromthe Pacific and Indian Oceans into the Atlantic Ocean at about the depth of the thermocline. Webelieve these return flows to the Atlantic Ocean occur through the Drake Passage between Antarcticaand South America and through the Bering Sea into the Arctic Ocean and thence to the AtlanticOcean. The return flows in the high latitudes of the North Atlantic sink as the NADW.

A. Because of this pattern of circulation, the deep water of the world’s oceans generally gets olderfrom the Atlantic Ocean to the Indian Ocean to the Pacific Ocean. The time the water stays out ofcontact with the atmosphere, that is, its residence time, increases toward the Pacific Ocean. Theresidence time of the bottom waters of the Atlantic Ocean is 200–500 years; that of the PacificOcean is 1,000–2,000 years.

B. By upwelling as described above.

C. This is a difficult question and one that is asked by a number of scientists today. The answer is ofconcern to the world’s people because of the link to climate. It is likely that any substantial melt-ing of sea ice and the continental glacier of Greenland would add fresh water to the surface of theocean in the high latitudes of the North Atlantic for some period of time. This would affect therate of deep water formation because of the change in the salt content of surface ocean water. Inturn, the pattern of the conveyor belt circulation of the ocean could be altered. One suggestion isthat there would be a less intense flow of warm water northward by the Gulf Stream, and the cli-mate of Europe would cool.

3. A. Because of the short residence time of CO in the atmosphere (about 70 days) and the fact that the upwelling of air near the equator effectively separates air exchange in the troposphere between the two hemispheres, the gas would not be evenly distributed. Higher concentrations would be found in the Northern Hemisphere troposphere than in the Southern Hemisphere. In fact, observations show a strong gradient in the concentration of CO between the two hemi-spheres, with higher concentrations in the North.

B. The barrier is the convergence of the North Equatorial Current and the South Equatorial Current near the equator and their westward flows. This converging pattern inhibits exchange of surface waters between the Northern Hemisphere and the Southern Hemisphere. Floating tar balls in the North Atlantic would be caught in the North Equatorial Current and transported westward to thenorthward-moving Gulf Stream.

C. No, the NADW moves southward at depth in the North Atlantic Ocean into the South Atlantic Ocean. The ABW moves north at depth. Neither current is involved with surface ocean currents.

4. The fine dust particles would take 4 km x 105 cm/km = 4 x 105 cm ÷ 500 cm/yr = 800 years. In actu-al fact, they settle much faster because they are encapsulated in the fecal pellets of animal plankton(zooplankton) in the ocean. During feeding the zooplankton inadvertently pass them through theirguts and excrete them contained in bigger mucilaginous fecal particles that sink at rates of 350 m/day.

5. To answer this question requires an understanding of some basic chemistry. The equilibriumbetween a gas and a solution is normally given by Henry’s Law, which states that the concentrationof the gas in the solution equals a constant (known as the Henry’s Law constant) times the partialpressure of the gas. For CO2, the expression is

[CO2] = KHPCO2


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The brackets around CO2 denote its concentration in moles/L. The partial pressure of CO2 (PCO2) is in atmospheres. The value of KH varies with the composition of the solution and the temperature.For seawater at 25°C, KH equals 10-1.54.

A. Carbon dioxide dissolves in seawater to an extent determined by the CO2 concentration in theatmosphere and the reactions that occur in the seawater. At 25°C and with an atmospheric CO2concentration of PCO2 = 10-3.45 atm, the relationship is

[CO2] = 10-1.54 x 10-3.45 = 10-4.99 mole/L

B. In 1700 for a PCO2 = 10-3.55 atm, we have

[CO2] = 10-1.54 x 10-3.55 = 10-5.09 mole/L

C. The percentage change is [(10-4.99 mole/L – 10-5.09 mole/L) ÷ 10-5.09 mole/L] x 100 = 26%. Thus,the dissolved CO2 concentration in the surface ocean has changed by this percentage over thepast 300 years because of fossil fuel combustion and biomass burning.

6. A. The percentage is [(0.9 x 10-4 cal/cm2/min + 0.9 x 10-5 cal/cm2/min) ÷ 0.5 cal/cm2/min + 0.9 x 10-4 cal/cm2/min + 0.9 x 10-5 cal/cm2/min] x 100 = 0.02%.

B. The amount of energy reaching the earth’s surface and absorbed is 343 W/m2 x 0.49 = 168 W/m2.This energy is used to heat the atmosphere and surface of the earth; to evaporate water; to gener-ate rising air masses (thermals); to drive wind, waves, and currents; and for photosynthesis.

C. The energy is reabsorbed, keeping the earth warm.

7. A. Provided there were no processes restoring the continents, they would be reduced to sea level by erosion in 150 x 1012 m2 x 840 m = 126 x 1015 m3 x 106 cm3/m3 = 126 x 1021 cm3 x 2.7 g/cm3 = 340 x 1021 g ÷ 200 x 1014 g/yr = 17 x 106 years.

B. This is a question that has plagued geologists for two centuries. There must be processes that addmass to the continents to keep them above sea level. Certainly lavas originating in the interior ofthe earth add material to the continents. In subduction zones, not all the rock of the oceanic crustis transported down toward the interior of the earth. Some of it is added to the continents andincreases their area and thickness. Finally, continents often collide during their movement aboutthe earth’s surface. In the collisions, the continents act as a great vise, squeezing sediments origi-nally derived from their erosion and other sources into high mountain ranges. This action addsvolume back to the continents. The collision of India with Asia followed by the formation of theHimalayan Mountains is an example of such an event. Thus, there is a great rock cycle at work inwhich the continents are eroded and their materials deposited in the ocean. The sediments of theoceans are buried to great depths or transported to subduction zones. In either case, the sedimen-tary material is eventually returned to the continents to be uplifted to their surfaces and eroded,completing the cycle.

8. A. This question simply uses the concept of residence time, where λ = mass ÷ flux; therefore, flux = mass ÷ λ = 1.8 x 1018 metric tons ÷ 600 x 106 yr = 3 x 109 metric tons/yr = 30 x 1014 g/yr. Interestingly, this flux is much less than that of erosion and thus deposition in the oceans today. This implies that today is somewhat unusual in terms of geologic history. We know that to be the case from other geological information.


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B. Sediment (clay, mud, silt, sand, gravel, and skeletal and organic materials) enters the sedimentaryrock reservoir by erosion of rocks, transportation of the eroded debris, and subsequent depositionon the seafloor, followed by burial in some cases to depths of 12 km. The sediment “gets out” byuplift due to plate tectonic movements and re-erosion of the sedimentary mass.

9. The chemical species in anoxic soils are chiefly reduced forms of the compounds CH4, CO, NH3, and H2S(see Figures 9, 10, 15, and 19). CO2 occurs as well, although it is a relatively oxidized form of carbon. In awell-aerated soil, we might expect O2, CO2, N2, and SO2.

10. Your answer to Question 2 in the Chemistry section shows that the weathering of silicate mineralssubtracts CO2 from the atmosphere. This subtraction must be balanced by an addition of CO2 fromother processes or the atmosphere would run out of CO2 in about 6,000 years. Those processesinvolve hydrothermal reactions at midocean ridges and the subduction and/or burial of carbonateminerals to realms of higher pressure and temperature, where carbonates are converted to silicates,and CO2 is released back to the atmosphere by volcanism and other processes.

Ecology

1. This question is designed to give the student a feeling for some of the characteristics of two impor-tant ecosystems that are being severely impacted by human activities. These activities are the defor-estation of tropical rain forests and their conversion to pasture and urban areas and the eutrophica-tion of coastal marine environments.

A. The total net primary production of tropical rain forests and estuaries is:

Rain forests: 17 x 1012 m2 x 2,000 g dry matter/m2/yr = 34 x 1015 g dry matter x 0.45 = 15.3 x 1015 g C.Dividing by 106 g/metric ton, we get 34 x 109 metric tons dry matter and 15.3 x 109 metric tons C.

Estuaries: 1.4 x 1012 m2 x 1,800 g dry matter/m2/yr = 2.52 x 1015 g dry matter x 0.45 = 1.13 x 1015 g C.Dividing by 106 g/metric ton, we get 2.52 x 109 metric tons dry matter and 1.13 x 109 metric tons C.Notice that although the NPP of tropical rain forests and estuaries is similar, the smaller area of estu-aries leads to more than an order of magnitude difference between the total net primary productionof the two ecosystems.

B. The total plant mass (biomass) of these ecosystems is:

Rain forests: 17 x 1012 m2 x 20 kg C/m2 = 340 x 1012 kg C ÷ 103 kg/metric ton = 340 x 109 metrictons C ÷ 0.45 = 755 x 109 metric tons dry matter.

Estuaries: 1.4 x 1012 m2 x 0.45 kg C/m2 = 630 x 109 kg C ÷ 103 kg/metric ton = 6.3 x 108 metric tons C ÷ 0.45 = 1.4 x 109 metric tons dry matter. Notice the orders-of-magnitude difference between the biomass of tropical rain forests and estuaries.

C. In ten years, 9% of the area of rain forests was lost; this represents 755 x 109 metric tons dry matter x 0.09 = 68 x 109 metric tons dry matter x 0.45 = 30.6 x 109 metric tons C.

D. The grams of CO2 emitted to the atmosphere in this ten-year period would be 30.6 x 109 metrictons C x 106 g/metric ton = 30.6 x 1015 g C x (44 ÷ 12) = 112 x 1015 g CO2. This flux represents(Figure 11) the size of the atmospheric CO2 reservoir of 744 x 1015 g C x (44 ÷ 12) = 2,730 x 1015 gCO2. 112 x 1015 g CO2 ÷ 2,730 x 1015 g CO2 = 0.041, or 4% of the atmospheric reservoir.


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The flux would represent about 3 x 1015 g C/yr for the decade of the 1980s. The magnitude of theflux is too high, meaning that all the woody plant material was not burned to CO2. Some remainson the ground as waste from the cutting, and some has gone into lumber. The flux of CO2 to theatmosphere from all land-use activities in tropical rain forests for the 1980s was on the order of1–1.6 x 1015 g C/yr.

E. At a C:P ratio of 106:1, the additional plant productivity supported by the pollutant P would be 1.5 x 1012 g P/yr ÷ 31 g/mole = 48.4 x 109 moles P/yr x (106 ÷ 1) = 5.13 x 1012 moles C/yr. Divided by the estuary area of 1.4 x 1012 m2 = 3.7 moles C/m2/yr = 44 g C/m2/yr ÷ 0.45 = 98 g dry matter/m2/yr, or about a 5% increase in the productivity of the world’s estuaries.

F. In fact, because the pollutant P will be used more than once in plant productivity in estuaries, thetotal organic carbon that could be buried in the sediments of estuaries amounts to 140 x 1012 gC/yr (Figure 18). This flux does qualify as a negative biological feedback because the additional Phas been added to the earth’s surface by the human activities of fertilizer application to croplandsand sewage discharge. Some of this P makes its way to lakes and coastal marine environmentsand increases the productivity of such aquatic systems. The flux represents (140 x 1012 g C/yr ÷ 6x 1015 g C/yr) x 100 = 2% of the fossil fuel flux in 1995. Not much!

2. A. The total production of dry matter is 510 x 1012 m2 x 0.37 kg dry matter/m2/yr = 188.7 x 1012 kg dry matter/yr.

B. 188.7 x 1012 kg dry matter/yr x 0.45 = 84.9 x 1012 kg C/yr.

C. Let total production in the ocean = X; then production on land = 2X. Thus X + 2X = 188.7 x 1012

kg dry matter/yr, i.e., 3X = 188.7 x 1012 kg dry matter/yr; X = ocean production = 62.9 x 1012 kgdry matter/yr and 2X = land production = 125.8 x 1012 kg dry matter/yr.

3. Autotrophy is the biochemical pathway by which an organism uses CO2 as a source of carbon andsimple inorganic nutrient compounds of N and P for synthesis of organic matter. In heterotrophy,more complex organic materials are used as the source of carbon for metabolic processes.

4. Prokaryotes—the Kingdom Monera, including the bacteria and cyanobacteria—take part in a varietyof biogeochemical processes (see Table 1). We often forget the fact that the bacteria are responsiblefor the decay of organic matter both on land and in the ocean. In other words, it is their metabolicactivity that returns CO2 and other gases and nutrients back to the environment. The cyanobacteriaare photoautotrophic and produce oxygen as a metabolic byproduct. These organisms were responsi-ble for the initial growth of oxygen in the earth’s atmosphere. The processes include CO2 fixation,nitrogen fixation, and oxidation of sulfur as examples.

5. Eutrophication is the set of processes leading to the overnourishment of a lake, river, or marine envi-ronment; consequent rapid plant growth and death; and oxygen deficiency of the system. This is anatural set of processes in certain environments. When it occurs because the excess nutrients comefrom fertilizers, sewage, detergents, etc., it is called cultural eutrophication. This term distinguishesthe natural situation from that produced by the activities of people.


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Biogeochemical Cycles

1. See Figure 22.

A. The residence time = λ = mass ofwater in the ocean divided by theflux of water to the ocean by rivers:1,370 x 106 km3 ÷ 40 x 103 km3/yr =34,250 years.

B. λ of Ca in the ocean is (1,370 x 106

km3 x 1012 L/km3 x 400 mg/L) ÷ (40x 103 km3/yr x 1012 L/km3 x 15mg/L) = 913,000 years.

2. Here we apply the concept of residence time in the context of how long it takes for a lake to recoverfrom a single input of a chemical into it. This is an environmental problem often encountered indeveloping, as well as developed, countries.

A. The volume of the lake is 5 km x 2 km x 0.1 km = 1 km3 x 1012 L/km3 = 1 x 1012 L.

B. The mass of mercury (Hg) is 1 x 1012 L x 1 x 10-6 g Hg/L = 1 x 106 g Hg.

C. The flux of mercury by the river to the lake is 2 x 1012 L/yr x 5 x 10-7 g Hg/L = 1 x 106 g Hg/yr;thus the residence time of Hg (λHg) in the lake is 1 x 106 g Hg ÷ 1 x 106 g Hg/yr = 1 year.

D. The time to reach a new equilibrium—that is, for the Hg input to work its way through the lake—is 3 x λHg = 3 x 1 yr = 3 yr. In this period of time, 95% of the Hg input would be gone.

3. This question provides the student with a feeling for the problem of “looking for a needle in ahaystack” that scientists have when observing the change in the size of the terrestrial living biomassdue to uptake and storage of anthropogenic CO2.

A. 600 x 109 tons C x 106 g/ton = 600 x 1015 g C. 135 x 1012 moles C/yr x 12 g/mole = 16.2 x 1014 gC/yr. The annual % change would be (16.2 x 1014 g C/yr ÷ 600 x 1015 g C) x 100 = 0.27% per year.

B. Most unlikely. For the decade of the 1980s, this would be only a 2.7% change in the mass of livingbiomass globally. This amount of change would be difficult to measure by field studies, even ifthey were aimed at seeing a change in the amount of organic carbon stored in terrestrial vegeta-tion. However, if the storage continues, such measurements could provide information in a fewyears.

4. A. The total DIC in the mixed layer is 360 x 1012 m2 x 102 m = 360 x 1014 m3 x 106 cm3/m3 = 360 x 1020 cm3 x 1.027 g/cm3 = 37 x 1021 g ÷ 103 g/kg = 37 x 1018 kg seawater x 2.2 x 10-3 moles C/kg seawater = 81.4 x 1015 moles C.

B. 2 x 109 tons C/yr = 2 x 1015 g/yr ÷ 12 g/mole = 166.7 x 1012 moles C/yr ÷ 37 x 1018 kg seawater = 4.5 x10-6 moles C/kg seawater/yr = 4.5 micromoles C/kg seawater/yr. Actual measurements of the changein the DIC content of seawater over time show an increase of about 1 micromole per year. The differ-ence is due to the fact that the anthropogenic carbon taken up by the ocean mixes deeper in the oceanthan the average depth of the mixed layer used in the problem, on average about 300–400 meters.

Plants Animals

Soil

LAND ATMOSPHERE

Water Biota

Sediment

OCEAN

Figure 22.


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C. Using 1 micromole per year as the average carbon uptake since 1975, we get 1 x 10-6 mole C/kg/yrx 21 yr = 21 x 10-6 moles C/kg ÷ 2.2 x 10-3 moles C/kg x 100 = 0.95%. It was not until the late 1980sthat scientists were able to measure these small changes in seawater DIC accurately and precisely.

5. Referring to Table 3, the total amount of water in the atmosphere as water vapor is 0.013 x 106 km3; thusthe residence time of water vapor in the atmosphere calculated with respect to total evaporation (mustequal precipitation) is 0.013 x 106 km3 H2O ÷ 496 x 103 km3 H2O/yr = 0.026 yr x 365 days/yr = 9.6 days.

B. No, the pollutant would not mix evenly throughout the troposphere because of water’s very shortresidence time in the atmosphere. In general, the dust and aerosol content of the troposphereexhibits a regional pattern and is most concentrated near sources, downwind from sources, and inregions of relatively dry climate. However, the dust plume in the troposphere derived from theSahara and Sahel areas of Africa can be seen in satellite images extending all the way across theAtlantic Ocean.

6. The three major processes are (1) inputs of volcanic CO2 derived from the metamorphism of CaCO3 toCaSiO3, (2) weathering of silicate minerals like CaSiO3 and then deposition of calcium carbonate andsilica in the ocean, and (3) the evolution of plants and their effect on weathering. The reactions are

1. CaCO3 + SiO2 ⇒ CaSiO3 + CO2

2. CaSiO3 + 2CO2 + 3H2O ⇒ Ca2+ + 2HCO3- + H4SiO4

0, and then Ca2+ + 2HCO3

- + H4SiO40 ⇒ CaCO3 + SiO2 + 3H2O + CO2

3. CO2 + H2O ⇔ CH2O + O2

7. The connection is simply the fact that when organic matter is buried in sediments of the ocean, oxy-gen not used to oxidize the organic matter is left in the atmosphere. With the continuous burial oforganic matter on the seafloor, the oxygen content would increase in a few millions of years to levelsthat would lead to burning of forests and grasslands. This does not happen because the buried organ-ic matter is returned to the earth’s surface through uplift by plate tectonic processes. On exposure tothe atmosphere, the organic matter is oxidized, and the oxygen is removed from the atmosphere.

8. The major difference is that there is not a major gas of phosphorus that resides in the atmosphere oris transported through it (see Figure 18). This statement is not true of carbon, nitrogen, and sulfur.All of these elements have important gaseous compounds in the atmosphere that exchange with theearth’s surface. (See the text sections on the cycles of C, N, and S.)

9. DMS in the troposphere reacts fairly rapidly with hydroxyl radical in the presence of light, water,and oxygen to make sulfate aerosol. On the contrary, OCS is inert in the troposphere but is convertedin the stratosphere to sulfate aerosol (see Figure 19).

10. The difference in residence time implies that ammonia reacts more rapidly than carbon monoxide inthe atmosphere (see Figures 10 and 16). This is simply a consequence of the fact that the shorter theresidence time (lifetime) of a chemical compound in a reservoir, the more reactive the substance.


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11. This question and the following five questions relate to relationships depicted in the figures of thebiogenic trace gases. They are designed to help the student to study the figures and interpret them.

The major reaction in the atmosphere coupling the biogeochemical cycles of CH4, CO, and CO2 is theoxidation of the reduced carbon gas CH4 to CO and then on to CO2 by OH*. About 60% of theenhanced greenhouse forcing is due to CO2, and about 20% is due to CH4. The major anthropogenicsources of emissions of CH4 to the atmosphere are fossil fuel burning and leakage from gas transmis-sion pipelines, biomass burning, landfills, rice paddies, and enteric fermentation in domesticatedanimals. CO emissions come from fossil fuel and biomass burning. Land-use activities and fossil fuelburning are the major anthropogenic sources of CO2 to the atmosphere. A doubling of atmosphericCO2 concentration could lead to a 2.5°C increase in mean global temperature. Molecule for molecule,CH4 is about 20 times more effective as a greenhouse gas than CO2. (Compare the amount of tem-perature change per ppbv for the two gases.)

12. The ratio is (42 + 20 + 78) x 106 metric tons N/yr ÷ 126 x 106 metric tons N/yr = 1.1:1. The anthro-pogenic nitrogen fluxes on land slightly exceed the natural biological fixation flux! The minimumpercentage is (21 x 106 metric tons N/yr ÷ 140 x 106 metric tons N/yr) x 100 = 15%. The additionalnitrogen flux to the ocean is a cause of eutrophication of coastal marine environments.

13. The ratio is 206 x 106 metric tons N/yr ÷ 62 x 106 metric tons N/yr = 3.3:1. This is a difficult questionto answer. On the time scale of a century, it is very likely that the flux of N to the ocean from riversand groundwaters will increase because of continuous use of industrial fertilizers, atmosphericdeposition of anthropogenic nitrogen, and disposal of sewage. If there is climatic change on this timescale, it is uncertain whether the upwelling rate of the world’s oceans will change. However, awarming of the global surface layer of the ocean would most likely lead to a slowing of upwellingand thus delivery of nutrients to the surface ocean.

14. The range is 143 to 339 years. 73 x 106 tons N ÷ (2.9 + 0.1 + 0.02 + 0.01 + 1.4) x 106 tons N/yr = 339years. 1,500 x 106 tons N ÷ (5.2 + 0.3 + 0.2 + 2.2 + 2.6) = 143 years. N2O accounts for about 9% of theenhanced greenhouse effect.

15. The percentage from human activities is [(21 + 3) x 106 tons N/yr ÷ (21 + 3 + 20) x 106 tons N/yr] x100 = 54.5%. Human activities substantially interfere with the fluxes of NOx. Environmental prob-lems associated with the anthropogenic fluxes are acid deposition, photochemical smog, andincreased tropospheric ozone, a greenhouse gas. The principal reaction leading to destruction of NOxin the atmosphere is photochemical. The products of the reaction are nitric acid, peroxylacetylnitrate, and organic nitrates. The reaction is NO2 + OH*+ light ⇒ HNO3.

16. The ratio is 1,085 x 106 tons P/yr ÷ 186 x 106 tons P/yr = 5.8:1.

The residence time of P in the marine biota relative to the recycling flux is 73 x 106 tons P ÷ 1,085 x106

tons P/yr = 0.07 years. That for P in the land biota is 1,800 x 106 tons P ÷ 186 x 106 tons P/yr = 9.6 years.The phytoplanktyon of the ocean “turn over” much more rapidly than do terrestrial plants. There aremore generations of death and birth for marine plankton than for most plants growing on land.

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Acid deposition—the fallout of acidic substances, primarily of nitrogen and sulfur, from the atmosphereon to the earth’s surface in rain, snow, or other forms. Acid rain has a pH generally less than 5 (aver-aged over a year).

Aerosol—the suspension of very fine, generally micrometer-sized, solid and liquid particles in the atmo-sphere.

Albedo—the amount of incident radiation that is reflected by a surface and thus does not contribute tothe heating of the surface. The albedo of the whole earth is approximately 30%. The albedo of cleansnow is about 90% and that of water is about 10%.

Anaerobic—an organism that does not need oxygen to carry on its metabolism or an environment with-out oxygen.

Anoxic—without oxygen.

Anthropogenic—of, relating to, or influenced by the impact of humans on nature.

Atom—the smallest component of an element having the chemical properties of the element. An atomconsists of a nucleus of neutrons and protons and one or more electrons bound to the nucleus by electri-cal attraction.

Autotrophy—the biochemical pathway by which an organism uses carbon dioxide as a source of carbonand simple nutrient compounds for synthesis of organic matter.

Autotrophic system—an environment in which the difference between gross photosynthesis and grossrespiration is positive. In such a terrestrial or aquatic environment, the net transfer of carbon dioxide isinto the system.

Bacteria—one-celled organisms having a spherical, spiral, or rod shape belonging to the Kingdom Monera.

Benthic—of, relating to, or occurring at the bottom of a body of water.

Bioessential—required by virtually all living organisms. The major bioessential elements are oxygen,carbon, nitrogen, phosphorus, sulfur, potassium, magnesium, and calcium. Minor or trace quantities ofiron, manganese, copper, zinc, boron, silicon, molybdenum, chlorine, vanadium, cobalt, and sodium arealso required by organisms.

Biogenic gas—a gas whose production or consumption on earth is accomplished by biological reactions.

Biogeochemical cycle—representation of biological, geological, and chemical processes that involve themovement of an element or compound about the surface of the earth.

Biogeochemical system—the interactive system of biogeochemical processes and cycles of elements andcompounds.

Biogeochemistry—the discipline that links various aspects of biology, geology, and chemistry to investi-gate the surface environment of the earth.


Glossary


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Biological productivity—the rate of production per unit area of organic matter by producer organisms.For example, the rate may be given as grams of carbon per square meter per year for a marine grasscommunity. There are several kinds of productivity. Gross primary production (GPP) refers to the totalamount of plant material produced by photosynthesis in a defined area in an interval of time. Net pri-mary production (NPP) is the net amount of plant material produced per unit area per unit time and isthe difference between GPP and cell respiration. Net ecosystem production (NEP) is the differencebetween GPP and cell respiration plus heterotrophic processes of decay.

Biological pump—the set of processes by which organic carbon is exported from the surface ocean tothe deep sea.

Biomass—the amount of living matter in a unit area or volume of habitat. For example, the total bio-mass of the world’s tropical rain forests is 42 kilograms of dry matter per square meter of forest, or atotal of 420 billion tons of dry matter (equivalent to approximately 170 billion tons of carbon).

Biosphere—the living and dead organic components of the earth. Sometimes this term is used in thesame way as the term ecosphere in this module, and sometimes for only the living animals and plants.

Climate—the characteristic long-term environmental conditions of temperature, precipitation, winds,etc., in a region or for the globe at present or in the past (paleoclimate).

Cloud condensation nuclei—airborne particles of very small size, generally less than one micrometer indiameter, that serve as sites on which liquid cloud droplets condense when an air mass is supersaturat-ed with water vapor. The particles are commonly composed of water-soluble material.

Coccolithophoridae—a family of planktonic algae that build skeletons of micrometer-sized disc-shapedplates of calcite, called coccoliths.

Concentration—the fraction of the total of a substance made up of one component. For example, seawa-ter contains 400 parts per million by weight of calcium. Concentration is also expressed in moles perliter or kilogram or in percent (that is, parts per hundred), parts per thousand (°/°°), per million (ppm),per billion (ppb), and so forth, either by weight or by volume.

Coupled—the condition in which information from one part of the system is provided to, and influencesthe behavior of, other parts. The biogeochemical cycles of the elements necessary for life are coupledthrough processes that are essential for life, e.g., photosynthesis and respiration.

Crust—the outer layer of the earth, enriched in silicon, sodium, and potassium and having a thicknessof 35 kilometers beneath the continents and 10 kilometers beneath the oceans.

Cryosphere—the icy part of the earth; its continental and mountain glaciers, ice sheets, and ice shelves;a reservoir in the earth’s surface system.

Decay—the oxidative process of conversion of organic tissue to simpler organic and inorganic com-pounds. The oxidizing agent may be diatomic oxygen (O2), nitrate (NO3

-), or other chemical compounds.

Denitrification—the conversion, principally by bacteria, of compounds of nitrogen in soils and aquaticsystems to nitrogen gas (N2) and nitrous oxide gas (N2O) and the eventual release of these gases to the atmosphere.

Diagenesis—the collection of physical, chemical, and biological processes that operate on a sedimentafter deposition.

Diatom—planktonic and benthic freshwater and marine algae that commonly use silicon to build askeleton of opal.


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Dry deposition—the deposition of materials from the atmosphere on to the earth’s surface in the formof solid particles. Such particles also may be “washed out” of the atmosphere by rain.

Ecosphere—the system that includes the biosphere and its interactions with the air, water, and soils andsediments of the earth.

Ecosystem—complex of a community or group of communities and the environment that functions asan ecological unit in nature.

Electromagnetic spectrum—the entire range of radiation. The wavelengths (distances between adjacentpeaks) of the electromagnetic waves within the spectrum range from kilometers for radio waves to bil-lionths of a meter (nanometers) for X rays.

Entropy—a scientific measure of the degree of disorder in a system. The greater the disorder the greater isthe entropy of the system. The Second Law of Thermodynamics states that entropy is always increasing.

Equilibrium—stable, balanced state in which all influences on a system are countered by others.

Erosion—the set of processes by which the surface of the earth is worn away by the action of water,wind, glacial ice, etc.

Euphotic zone—the upper, lighted zone of the ocean or a lake in which most of the productivity ofplants occurs. In the ocean, the euphotic zone extends from the surface to a depth where the light inten-sity is reduced to about 0.1–1.0% of that available at the surface. The depth of the euphotic zonedepends on season and latitude.

Eutrophication—the set of processes leading to overnourishment of an aquatic system in nutrients,rapid plant growth and death, and oxygen consumption and deficiency in the system. These processesoccur naturally in some aquatic systems but may be speeded up by additions of nutrients from humanactivities (e.g., fertilizer application) to the systems. Human-induced eutrophication is often called cul-tural eutrophication.

Evaporation—the physical process by which water is converted from liquid to vapor and is transportedinto the atmosphere.

Evapotranspiration—the combined processes of evaporation and transpiration.

Evasion—the escape or release of a gas from the surface of the ocean or land to the atmosphere.

Evolution—the pattern of development and change in a variable from one state to another. Biologicalevolution describes the pattern of emergence, development, and extinction of organic species throughgeologic time.

Feedback—a process or mechanism in which some fraction of the output is returned or “fed back” tothe input. Feedback loops may either stabilize (negative feedback) or destabilize (positive feedback) asystem undergoing a perturbation. These feedback loops exist in both the biogeochemical cycles and theclimate system.

Fermentation—the bacterial process of conversion of sugars to carbon dioxide.

Fixation—see Nitrogen fixation.

Flux—the movement of a variable or a substance into or out of a reservoir.

Foraminifera—animal plankton (zooplankton) in the ocean belonging to the Phylum Protozoa that com-monly have a shell of calcium carbonate.


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Forcing—the ability of a variable, like the concentration of a greenhouse gas in the atmosphere, toinduce a change in a system. A forcing function controls the behavior of a system and often makes itregular and predictable.

General circulation model (GCM)—a simulation, usually performed on a large computer, of the large-scale, or general, wind and ocean systems on earth to calculate climate and its changes.

Geologic time scale—a calendar of earth history. The time scale is divided into variable time units ofeon, era, period, and epoch (see Figure 5).

Glacial stage—an extended cold interval of time within the Pleistocene Epoch in which continental glac-iers covered much of the Northern Hemisphere continents, atmospheric CO2 concentrations were low,and sea level was low.

Greenhouse effect—the warming of the earth’s atmosphere and surface by the atmospheric greenhousegases. These gases absorb and reradiate longwave radiation from the earth, keeping it in the atmosphereand thus warming the global temperature. Without the natural greenhouse effect, the planet would beabout 33°C cooler than its global mean annual temperature of 15°C, that is, –18°C. Because of inputsfrom human activities, these gases are increasing in concentration in the atmosphere. This may lead toan enhanced greenhouse effect and warming of the planet.

Greenhouse gas—an atmospheric gas that absorbs and radiates energy in the infrared part of the elec-tromagnetic spectrum. Such gases include water vapor, carbon dioxide, methane, nitrous oxide, tropo-spheric ozone, and the synthetic chlorofluorocarbon gases. These gases warm the atmosphere and theearth’s surface below.

Groundwater—the water beneath the ground, largely formed by the seepage of surface water downward.

Heterotrophic system—an environment in which the difference between gross photosynthesis and grossrespiration is negative. In such a terrestrial or aquatic environment, the net transfer of carbon dioxide isout of the system.

Heterotrophy—a biochemical pathway in which organic substrates are used by organisms to makeorganic matter.

Hothouse—an extended period of geologic time during which the earth was warm.

Hydrosphere—the watery envelope surrounding the earth; a reservoir in the earth’s surface system.

Hydrothermal reaction—a chemical reaction involving hot water and minerals in a rock.

Hydroxyl radical (OH*)—the excited chemical compound of hydrogen and oxygen in the atmospherewith an imbalance of electric charge. The hydroxyl radical is responsible for the oxidation of manychemically reduced gases emitted from the surface of the earth.

Ice age—a glacial stage, especially within the Pleistocene Epoch, beginning about 1.8 million years ago.

Ice house—an extended period of geologic time in which the earth was cool.

Infrared radiation—the region of the electromagnetic spectrum with wavelengths longer than visiblelight (about 1 micrometer) but shorter than microwaves (about 1 millimeter). Commonly known as heat.Radiation emitted from the earth back to space is predominantly infrared radiation.

Interglacial stage—an extended warm interval of time within the Pleistocene Epoch in which the conti-nental glaciers retreated and atmospheric CO2 concentrations and sea level were low.

Ion—an electrically charged atom or group of atoms formed by the loss or gain of one or more electrons.A positive ion, the cation, is created by an electron loss, and a negative ion, the anion, is created by anelectron gain.


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Irreversible process—a process in which the entropy change is greater than zero. After the process iscomplete, the system is more disordered than and different from its initial state.

Kerogen—fossil organic matter dispersed throughout a rock.

Leaching—the selective removal of substances from a substrate, usually with water. For example, rainwa-ter percolating through a soil can dissolve nitrogen and transport it to the groundwater. This is leaching.

Lifetime—a measure of the reactivity of an atmospheric chemical compound. The more reactive thecompound, the shorter its atmospheric lifetime. Analogous to residence time.

Limestone—a sedimentary rock consisting predominantly of calcium carbonate minerals.

Limiting nutrient—the chemical compound, generally inorganic, that limits productivity in a terrestrialor aquatic environment. Examples are nitrate, phosphate, and iron.

Lithosphere—the dynamic subdivision of earth on the order of 100 kilometers in thickness forming theouter, rigid part of the planet. Also, the solid portion of the earth, composed of minerals, rocks, andsoils; a reservoir in the earth’s surface system.

Mantle—the portion of earth between its crust and its innermost zone (the core). The mantle is enrichedin magnesium and iron and has a thickness of about 2,900 kilometers.

Metamorphism—the set of processes that lead to a change in the structure or composition of a rock dueto pressure and temperature. A metamorphic rock is formed from a preexisting rock by an increase inpressure and temperature.

Methanogenesis—the conversion of organic material to methane, principally by bacteria.

Methanotrophy—the conversion of methane to carbon dioxide, principally by bacteria.

Midocean ridge—any of several seismically active, submarine mountain ranges that are found in theAtlantic, Indian, and Pacific Oceans. These ridges are regions where the seafloor originates and are thesource of the lithospheric plates.

Mixotrophy—the use of both organic and inorganic materials to make organic matter.

Mole—one gram atomic weight of an element or one gram molecular weight of a compound. One gramatomic weight of an element is its atomic weight expressed in grams (i.e., the atomic weight of oxygenis 16; its gram atomic weight is 16 grams). One gram molecular weight of a compound is its molecularweight expressed in grams (i.e., the molecular weight of carbon dioxide is 44; its gram molecular weightis 44 grams).

Molecule—the smallest physical unit of an atom or compound, consisting of one or more similar atomsin an element and two or more different atoms in a compound.

Negative feedback—a process or mechanism that relieves or subtracts from an initial perturbation to a system.

Net primary production—see Biological productivity.

Nitrification—the conversion of ammonium to nitrite and nitrate by nitrifying bacteria.

Nitrogen fixation—the conversion of diatomic nitrogen gas (N2) to ammonium by bacteria. Also, the indus-trial conversion of free nitrogen into combined forms used as starting materials for fertilizers and explosives.

Nutrient—a substance that supplies nutrition to a living organism, like phosphorus and nitrogen.

Organic—pertaining to a class of chemical compounds that include carbon as a component; characteris-tic of or derived from living organisms.


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Oxidation—the removal of electrons from an atom or molecule. Oxidizing capacity is the intrinsic abili-ty of a system to oxidize reduced substances.

pH—the negative logarithm of the effective hydrogen ion concentration, used in expressing both acidityand alkalinity on a scale whose values run from 0 to 14, with 7 representing neutrality. Numbers lessthan 7 denote increasing acidity, and numbers greater than 7 increasing alkaline (basic) conditions.

Photoautotrophy—the conversion of inorganic carbon into organic matter in the presence of light.

Photochemical—pertaining to chemical reactions involving chemical compounds in the presence oflight. Urban smog is the result of a complex series of photochemical reactions involving ozone, nitrogenand sulfur oxides, and hydrocarbons.

Photolysis (photodissociation)—pertaining to chemical reactions triggered by light that convert a com-plex compound to more simple products. The photolysis of ammonia is an example: 2NH3 ⇒ N2 + 3H2.

Photosynthesis—the synthesis of complex organic materials (e.g., carbohydrates) from carbon dioxide, water,and nutrients, using sunlight as a source of energy with the aid of chlorophyll and associated pigments.

Phytoplankton—minute plant life that passively floats in a body of water. The phytoplankton are at thebase of the food chain in the ocean.

Plankton—minute plant and animal life of the ocean ranging in size from 5 micrometers to 3 centime-ters. The plant plankton are the phytoplankton; the animal plankton are the zooplankton.

Plate tectonics—the theory of global tectonics in which the lithosphere is divided into a number ofcrustal plates that move on the underlying plastic asthenosphere. These plates may collide with adjacentplates, slide under or over them, or move past them in a nearly horizontal direction. The sources of theplates are the great midocean ridges of the world’s oceans, where hot molten material upwells fromwithin the earth. The plates are destroyed at subduction zones, like that along the western margin of thePacific Ocean, where they sink down into the underlying asthenosphere.

Positive feedback—a process or mechanism that reinforces or adds to an initial perturbation of a system.

Precipitation—the removal of water from the atmosphere and its deposition on the earth’s surface in theform of rain, ice, or snow.

Prokaryote—any cellular organism that has no membrane about its nucleus and no organelles in thecytoplasm except ribosomes. Prokaryotic genetic material is in the form of single, continuous strandsforming coils or loops, characteristic of all organisms of the Kingdom Monera, such as bacteria orcyanobacteria.

Protozoan—eukaryotic organism of the Kingdom Protoctista, Phylum Protozoa, with a membrane-bound nucleus and organelles within a mass of protoplasm. Planktonic foraminifera and radiolarianswhich secrete shells of calcium carbonate and opal, respectively, are members of the group.

Radical—an electronically excited compound with an imbalance of electric charge, which enables it toreact rapidly with another molecule.

Redfield ratio—the relatively constant ratio of 106:16:1 of the bioessential elements carbon, nitrogen,and phosphorus in marine plankton. The concept of the Redfield ratio has been applied to the terrestrialrealm as well as to organic matter in soils and sediments.

Reduction—the chemical process by which an atom or a molecule gains electrons.


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Reservoir or stock—a part of a system that can store or accumulate and be a source of one of the sub-stances that compose the system. For example, the atmosphere is a reservoir of the surface system of theearth (the ecosphere). It can store water vapor released to it from the land by evapotranspiration andfrom the ocean by evaporation, and return the water to the earth’s surface as precipitation.

Residence time—the total mass of a substance in a reservoir divided by its rate of inflow or outflow.The residence time is a measure of the reactivity of the substance in the reservoir. For example, the resi-dence time of sodium in the ocean is very long (55 million years). Therefore, sodium does not enter intochemical or biochemical reactions that remove it very rapidly from the ocean. In contrast, the residencetime of dissolved silica in the ocean is about 20,000 years. This compound is readily taken up by certaintypes of plankton to build their skeletons.

Respiration—the physical and chemical processes by which an organism supplies its cells and tissues withthe oxygen needed for metabolism and releases carbon dioxide formed in the energy-producing reactions.

Reversible process—a process in which the change in entropy is zero. In general, after the process iscomplete, the state of the system is as it was initially.

Saturation—the degree to which a solution or a gas is at equilibrium with one of its components. It ismeasured in several different ways. For example, a humidity of 125% would be a supersaturation of 25%with respect to water vapor in the air. Saturation of seawater with respect to the mineral calcite (CaCO3)of 50% would mean that the seawater was 50% undersaturated with respect to calcite. If a lake watercontained exactly enough dissolved CO2 to be in equilibrium with the atmosphere, it would have a satu-ration of 100% with respect to CO2.

Sedimentary rock—a rock formed from the erosion of preexisting rocks and the deposition of the erod-ed materials as sediment. Sedimentary rocks are also formed by inorganic or biological precipitation ofminerals from natural waters.

Shortwave radiation—generally, the region of the electromagnetic spectrum with wavelengths shorter than0.5 micrometers. Solar radiation has an important component of shortwave radiation of varying intensity.

Solar radiation—the electromagnetic radiation emitted by the sun. It includes energy wavelengths fromthe very short ultraviolet (<0.2 micrometers) to about 3 micrometers.

Stratosphere—the region of the upper atmosphere extending upward from the troposphere to about 30kilometers above the earth’s surface. This region is characterized by an increase in temperature as alti-tude increases.

Subduction zone—the juncture of two lithospheric plates where the collision of the plates results in oneplate’s being drawn down or overridden by another plate. This region is the sink of the crustal plates ofthe earth.

System—a selected set of interactive components. An example of a simple system is an air conditioningunit. A biogeochemical system consists of reservoirs, processes and mechanisms, and associated fluxesinvolving material transport. The global climate system is very complex and involves all the physical,chemical, and biological interactions that control the long-term environmental conditions of the world.

Thermocline—the depth range in the ocean where the temperature decreases rapidly with increasingdepth. The thermocline is about one kilometer thick and extends from the base of the surface layer of theocean at a depth of 50–300 meters to a depth of about 800–1,000 meters.

Trace gas—a gas present in the atmosphere in a very low concentration (less than 1% of the composition ofthe atmosphere). For example, methane, nitrous oxide, and carbon monoxide are considered trace gases.

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Transitional phenomenon—a feature of a system that changes from one state to another. Such a changemay be relatively slow or more generally abrupt. In the boiling of water, the change from the state oflittle water motion to that of turbulence is a transitional phenomenon.

Transpiration—the process by which water in plants is excreted through a plant membrane as water vapor.

Troposphere—the lowest level of the atmosphere, up to 8–13 kilometers high, within which there is asteady drop in temperature with increasing altitude. It is the region where most cloud formations occurand weather conditions manifest themselves.

Ultraviolet radiation—the region of the electromagnetic spectrum with wavelengths longer than 0.5nanometers but shorter than 0.5 micrometers. Solar radiation has an important component of ultravioletradiation of varying intensity.

Uptake—generally, the incorporation of a substance into a solid or liquid. For example, the invasion ofCO2 into the ocean represents the uptake of CO2 from the atmosphere.

Upwelling—the upward movement of water from depths of typically 50–150 meters at speeds ofapproximately 1–3 meters per day. The upwelling of water generally results from the lateral movementof surface water. Upwelling zones in the ocean are found along the western margins of the continents, inequatorial regions, and at high latitudes of the Southern Hemisphere.

Vascular plant—either a plant with seeds that are not enclosed in a fruit or seed case, such as pine, fir,spruce, and other cone-bearing trees or shrubs (gymnosperm), or a flowering plant that producesencased seeds, such as oak, maple, and eucalyptus trees (angiosperm).

Volatilization—the conversion of a substance into the gas or vapor state and its emission into the environment.

Washout—the scavenging of particles from the atmosphere by rainfall and their subsequent depositionon the surface of the earth.

Weathering—the set of chemical, physical, and biological processes that lead to the disintegration ofminerals, kerogen, and rocks.

Wet deposition—the deposition on the earth’s surface of solid particles and dissolved chemical com-pounds in rain.

Zooplankton—minute animal life in a body of water that generally drift passively or swim very weakly.

Berner, E.K., and R.A. Berner, 1996, Global Environment: Water, Air, and Geochemical Cycles. Prentice Hall,Upper Saddle River, New Jersey.

Broecker, W.S., 1983, The ocean. Scientific American, vol. 1249, no. 3, 100–112.

Chiras, D.D., 1988, Environmental Science. Benjamin Cumming Publishing Co., Redwood City, California.

Ehrlich, P.R., A.H. Ehrlich, and J.P. Holden, 1977, Ecoscience: Population, Resources, Environment. W.H.Freeman and Co., San Francisco, California.

Garrels, R.M., F.T. Mackenzie, and C.H. Hunt, 1975, Chemical Cycles and the Global Environment: AssessingHuman Influences. William Kaufmann, Inc., Los Altos, California.

Graedel, T.F., and P.J. Crutzen, 1993, Atmospheric Change: An Earth System Perspective. W.H. Freeman andCo., San Francisco, California.

Gross, M.G., 1987, Oceanography: A View of the Earth, 4th ed. Prentice Hall, Englewood Cliffs, New Jersey.

Holland, H.D., and U. Patterson, 1995, Living Dangerously: The Earth, Its Resources, and the Environment.Princeton University Press, Princeton, New Jersey.

Mackenzie, F.T., Biogeochemistry. In Encyclopedia of Environmental Biology, vol. 1, Academic Press, Inc.,New York, 249–276.

Mackenzie, F.T., and J.A. Mackenzie, 1998 (2nd ed.), Our Changing Planet: An Introduction to Earth SystemScience and Global Change. Prentice Hall, Englewood Cliffs, New Jersey.

Schlesinger, W.H., 1991 (2nd ed.), Biogeochemistry: An Analysis of Global Change. Academic Press, SanDiego, California.

Skinner, B.J., and S.C. Porter, 1987, Physical Geology. John Wiley and Sons, New York.

Smil, V., 1997, Cycles of Life: Civilization and the Biosphere. Scientific American Library, New York.

Turekian, K.K., 1996, Global Environmental Change: Past, Present and Future. Prentice Hall, Upper SaddleRiver, New Jersey.

Wayne, R.P.,1991, Chemistry of Atmospheres. Oxford University Press, New York.

Supplementary Reading

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Andreae, M.O., 1987, The oceans as a source of biogenic gases. Oceanus, vol. 29, 27–35.

Berner, R.A., 1991, A model for atmospheric CO2 over geologic time. American Journal of Science, vol. 291,339–376.

Berner, R.A., and D.E. Canfield, 1989, A new model for atmospheric oxygen over Phanerozoic time.American Journal of Science, vol. 289, 333–361.

Christensen, J.W., 1991, Global Science. Kendall/Hunt Publishing Co., Dubuque, Iowa.

Guenther, A., C.N. Hewitt, D. Erickson, R. Fall, C. Geron, T. Graedel, P. Harley, L. Klinger, M. Lerdau,W.A. McKay, T. Pierce, B. Scholes, R. Steinbrecher, R. Tallamraju, J. Taylor, and P. Zimmerman, 1995,A global model of natural volatile organic compound emissions. Journal of Geophysical Research, vol.100, 8873–8892.

Houghton, J.T., G.J. Jenkins, and J.J. Ephraums (eds.), 1990, Climate Change: The IPCC ScientificAssessment. Cambridge University Press, Cambridge, U.K.

Houghton, J.T., L.G. Meira Filho, B.A. Callander, N. Harris, A. Kattenberg, and K. Maskell (eds.), 1996,Climate Change: The Science of Climate Change. Cambridge University Press, Cambridge, U.K.

Mantoura, R.F.C., J.-M. Martin, and R. Wollast (eds.), 1991, Ocean Margin Processes in Global Change.Wiley-Interscience, New York.

National Research Council, 1986, Global Change in the Geosphere-Biosphere. National Academy Press,Washington, D.C.

Skinner, B.J., and S.C. Porter, 1987, Physical Geology. John Wiley and Sons, New York.

Socolow, R., C. Andrews, F. Berkhout, and V. Thomas, 1994, Industrial Ecology and Global Change.Cambridge University Press, New York.

Stolz, J.F., D.B. Botkin, and M.N. Dastoor, 1989, The integral biosphere. In: Global Ecology, M.B. Rambler,L. Margulis, and R. Fester, eds. Academic Press, San Diego, California, 31–50.

Wollast, R., F.T. Mackenzie, and L. Chou, 1993, Interactions of C, N, P and S Biogeochemical Cycles andGlobal Change. Springer-Verlag, New York.

Woodwell, G.M., and F.T. Mackenzie, 1995, Biotic Feedbacks in the Global Climatic System. OxfordUniversity Press, New York.

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