Biogeochemistry: An Analysis of Global Change

By W.H. Schlesinger and Emily S. Bernhardt

Biogeochemistry—winner of a 2014 Textbook Excellence Award (Texty) from the Text and Academic Authors Association—considers how the basic chemical conditions of the Earth, from atmosphere to soil to seawater, have been and are being affected by the existence of life. Human activities in particular, from the rapid consumption of resources to the destruction of the rainforests and the expansion of smog-covered cities, are leading to rapid changes in the basic chemistry of the Earth.

This expansive text pulls together the numerous fields of study encompassed by biogeochemistry to analyze the increasing demands of the growing human population on limited resources and the resulting changes in the planet's chemical makeup.

The book helps students extrapolate small-scale examples to the global level, and also discusses the instrumentation being used by NASA and its role in studies of global change. With extensive cross-referencing of chapters, figures and tables, and an interdisciplinary coverage of the topic at hand, this updated edition provides an excellent framework for courses examining global change and environmental chemistry, and is also a useful self-study guide.

• Winner of a 2014 Texty Award from the Text and Academic Authors Association
• Calculates and compares the effects of industrial emissions, land clearing, agriculture, and rising population on Earth's chemistry
• Synthesizes the global cycles of carbon, nitrogen, phosphorous, and sulfur, and suggests the best current budgets for atmospheric gases such as ammonia, nitrous oxide, dimethyl sulfide, and carbonyl sulfide
• Includes an extensive review and up-to-date synthesis of the current literature on the Earth's biogeochemistry

• Language: English
• Publisher: Academic Press
• Release date: Dec 31, 2012
• ISBN: 9780123858757

W.H. Schlesinger

Dr. Schlesinger is one of the nation’s leading ecologists and earth scientists and a passionate advocate for translating science for lay audiences. A member of the National Academy of Sciences, he has served as dean of the Nicholas School of the Environment at Duke and president of the Cary Institute of Ecosystem Studies. He lives in Down East Maine and Durham, N.C. and continues to analyze the impacts of humans on the chemistry of our natural environment.

Book preview

Table of Contents

Cover image

Title page

Copyright

Dedication

Preface

Acknowledgments

Part I: Processes and Reactions

Chapter 1. Introduction

What is biogeochemistry?

Understanding the earth as a chemical system

Scales of endeavor

Lovelock’s gaia

Recommended Readings

Chapter 2. Origins

Introduction

Origins of the Elements

Origin of the solar system and the solid earth

Origin of the atmosphere and the oceans

Origin of life

Evolution of metabolic pathways

Comparative planetary history: earth, mars, and venus

Summary

Recommended Readings

Chapter 3. The Atmosphere

Introduction

Structure and Circulation

Atmospheric Composition

Biogeochemical Reactions in the Troposphere

Atmospheric Deposition

Biogeochemical Reactions in the Stratosphere

Models of the Atmosphere and Global Climate

Summary

Recommended Readings

Chapter 4. The Lithosphere

Introduction

Rock weathering

Soil chemical reactions

Soil development

Weathering rates

Summary

Recommended Readings

Chapter 5. The Biosphere: The Carbon Cycle of Terrestrial Ecosystems

Introduction

Photosynthesis

Respiration

Net Primary Production

Net Ecosystem Production and Eddy-Covariance Studies

The Fate of Net Primary Production

Remote Sensing of Primary Production and Biomass

Global Estimates of Net Primary Production and Biomass

Net Primary Production and Global Change

Detritus

Soil Organic Matter and Global Change

Summary

Recommended Readings

Chapter 6. The Biosphere: Biogeochemical Cycling on Land

Introduction

Biogeochemical cycling in land plants

Nutrient allocations and cycling in land vegetation

Biogeochemical cycling in the soil

Calculating landscape mass balance

Human impacts on terrestrial biogeochemistry

Summary

Recommended Readings

Chapter 7. Wetland Ecosystems

Introduction

Types of wetlands

Productivity in wetland ecosystems

Organic matter storage in wetlands

Microbial metabolism in saturated sediments

Anaerobic metabolic pathways

Wetlands and water quality

Wetlands and global change

Summary

Recommended Readings

Chapter 8. Inland Waters

Introduction

Lakes

Rivers

Estuaries

Human impacts on inland waters

SUMMARY

Recommended Readings

Chapter 9. The Oceans

Introduction

Ocean circulation

The composition of seawater

Net primary production

Sediment diagenesis

The biological pump: a model of carbon cycling in the ocean

Nutrient cycling in the ocean

Biogeochemistry of hydrothermal vent communities

The marine sulfur cycle

The sedimentary record of biogeochemistry

Summary

Recommended Readings

Part II: Global Cycles

Chapter 10. The Global Water Cycle

Introduction

The global water cycle

Models of the hydrologic cycle

The history of the water cycle

The water cycle and climate change

Summary

Recommended Readings

Chapter 11. The Global Carbon Cycle

Introduction

The modern carbon cycle

Temporal perspectives on the carbon cycle

Atmospheric methane

Carbon monoxide

Synthesis: linking the carbon and oxygen cycles

Summary

Recommended Readings

Chapter 12. The Global Cycles of Nitrogen and Phosphorus

Introduction

The global nitrogen cycle

Temporal variations in the global nitrogen cycle

Nitrous oxide

The global phosphorus cycle

Linking global biogeochemical cycles

Summary

Recommended Readings

Chapter 13. The Global Cycles of Sulfur and Mercury

Introduction

The global sulfur cycle

The global mercury cycle

Summary

Recommended Reading

Chapter 14. Perspectives

Recommended Reading

References

Index

Copyright

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Notices

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Library of Congress Cataloging-in-Publication Data

Schlesinger, William H.

 Biogeochemistry : an analysis of global change / William H. Schlesinger,

Emily S. Bernhardt. — 3rd edition.

   pages cm

 Includes bibliographical references and index.

 ISBN 978-0-12-385874-0

1. Biogeochemistry. I. Bernhardt, Emily S. II. Title.

 QH343.7.S35 2012

 572--dc23

2012030107

British Library Cataloguing-in-Publication Data

A catalogue record for this book is available from the British Library.

For information on all Academic Press publications

visit our website at http://store.elsevier.com

Printed in China

13 14 15 16 17 10 9 8 7 6 5 4 3 2 1

Dedication

To Lisa and Justin, for supporting our obsession with biogeochemistry and to Hannah and Gwyneth, with hopes for a more sustainable future.

Preface

This is a textbook about the chemistry of the surface of the Earth—the arena of life that is now increasingly affected by human activities. It is remarkable how many of our current environmental problems, ranging from climate change to ocean acidification, stem from changes in the Earth’s chemistry wrought by humans. We hope that students using this book will come to appreciate how the Earth functions naturally as a chemical system, what events have caused changes in Earth’s surface chemistry in the past, and what is causing our planet to change rapidly today. The book melds a wide range of disciplines from astrophysics to molecular biology and scales of time, from the origin of the Earth to the coming decades.

Like its previous editions, this book’s organization follows the structure of a class in biogeochemistry that we have taught for many years at Duke University—first by Schlesinger, then by both of us, and more recently by Bernhardt. Following our class syllabus, we have organized the book into two sections: Part I covers the microbial and chemical reactions that occur in the atmosphere, on land, in freshwaters, and in the sea. Part II is a set of shorter chapters that link the mechanistic understanding of the earlier chapters to a large-scale, synthetic view of global biogeochemical cycles.

Every section of this book has been revised from the most recent edition (1997), with special emphasis given to expand our treatment of freshwaters and aquatic ecosystems, to include satellite and model-derived global maps of Earth’s chemical characteristics, and to provide new global budgets for the major biogeochemical elements and mercury.

Throughout this book, we give special emphasis to the chemical reactions that link the elements that are important to life. The coupling of element cycling begins in biochemistry and constrains element cycling at the global level. In several locations, we show how computer models can be used to help understand and predict elemental cycling and ecosystem function. Many of these models are based on biochemistry and interactions among the biochemical elements. The models are useful in extrapolating small-scale observations to the global level. The models are often validated by observations taken from satellites, particularly since the deployment of NASA’s Earth Observing System (EOS). We hope that this book will link disparate fields ranging from microbiology to global change ecology—all of which are now part of the science of biogeochemistry.

This text provides the framework for a class in biogeochemistry. It is meant to be supplemented by readings from the current literature, so that areas of specific interest or recent progress can be explored in more detail. Although not encyclopedic, the text includes more than 4500 references to aid students and others who wish to explore any topic more thoroughly. Reflecting the book’s interdisciplinary nature, we have made a special effort to provide abundant cross-referencing of chapters, figures, and tables throughout.

As with this book’s first and second editions, we hope that this one will stimulate a new generation of students to address the science and policy of global change. Denial is not an option; our planet needs attentive stewardship.

WHS and ESB

Millbrook, NY, and Durham, NC

Acknowledgments

A huge number of people have helped to produce this volume, both directly through their efforts to supply us with data, references, reviews, and figures, and indirectly through their early influence on our careers. The latter include (for WHS) Jim Eicher, Joe Chadbourne, John Baker, Russ Hansen, Bill Reiners, Noye Johnson, Bob Reynolds, and Peter Marks; and (for ESB) Gene Likens, Barbara Peckarsky, Lars Hedin, Alex Flecker, Margaret Palmer, Pat Mulholland, and Bob Hall.

A number of friends helped immensely in the preparation of this edition by providing reviews of early drafts of our efforts, including Ron Kiene, Susan Lozier, Elise Pendall, Emma Rosi-Marshall, Dan Richter, Lisa Dellwo Schlesinger, Jim Siedow, Dave Stevenson, Mike Tice and Paul Wennberg. Several of them tested the drafts of certain chapters in their classes. Greg Okin and Daniel Giammar helped test the problem sets.

For this edition, we thank the following for references, figures, and comments—large and small—who have helped to produce a better book: Geoff Abers, Andy Andreae, Alison Appling, Dennis Baldocchi, Mike Behrenfeld, Neil Bettez, Jim Brown, Amy Burgin, Oliver Chadwick, Terry Chapin, Ben Colman, Jim Clark, Jon Cole, Bruce Corliss, Randy Dahlgren, Paul Falkowski, Ian Faloona, Jack Fishman, Jacqueline Flückiger, Wendy Freeman, Jim Galloway, Nicholas Gruber, Jim Hansen, Kris Havstad, James Heffernan, Ashley Helton, Kirsten Hofmockel, Ben Houlton, Dan Jacob, Steve Jasinski, Jason Kaye, Gabriel Katul, Ralph Keeling, Emily Klein, George Kling, Jean Knops, Arancha Lana, Steve Leavitt, Lance Lesack, Gene Likens, Gary Lovett, George Luther, Brian Lutz, John Magnuson, Pat Megonigal, Patrick Mitchell, Scott Morford, Karl Niklas, Ram Oren, Steve Piper, Jim Randerson, Sasha Reed, Bill Reiners, Joan Riera, Phil Robertson, Jorge Sarmiento, Noelle Selin, Gus Shaver, Hank Shugart, John Simon, Emily Stanley, Phil Taylor, Eileen Thorsos, Kevin Trenberth, Remco van den Bos, Peter Vitousek, Mark Walbridge, Matt Wallenstein, Kathie Weathers, and Charlie Yocum. In addition, Deb Fargione helped to keep more than 4500 references easily accessible.

We would also like to thank the many students of Duke’s Biogeochemistry course over the years for helping us refine our own understanding of biogeochemistry.

As always, any errors are our own, and we welcome hearing from you with comments, both large and small. You can contact us at [e-mail] or [e-mail].

Part I

Processes and Reactions

Chapter 1 Introduction

Chapter 2 Origins

Chapter 3 The Atmosphere

Chapter 4 The Lithosphere

Chapter 5 The Biosphere

Chapter 6 The Biosphere

Chapter 7 Wetland Ecosystems

Chapter 8 Inland Waters

Chapter 9 The Oceans

Chapter 1

Introduction

Outline

What Is Biogeochemistry?

Understanding the Earth as a Chemical System

Scales of Endeavor

Thermodynamics

Stoichiometry

Large-Scale Experiments

Models

Lovelock’s Gaia

What is biogeochemistry?

Today life is found from the deepest ocean trenches to the heights of the atmosphere above Mt. Everest; from the hottest and driest deserts in Chile to the coldest snows of Antarctica; from acid mine drainage in California, with pH < 1.0, to alkaline groundwaters in South Africa. More than 3.5 billion years of life on Earth has allowed the evolutionary process to fill nearly all habitats with species, large and small. And collectively these species have left their mark on the environment in the form of waste products, byproducts, and their own dead remains. Look into any shovel of soil and you will see organic materials that are evidence of life—a sharp contrast to what we see on the barren surface of Mars. Any laboratory sample of the atmosphere will contain nearly 21% oxygen, an unusually high concentration given that the Earth harbors lots of organic materials, such as wood, that are readily consumed by fire. All evidence suggests that the oxygen in Earth’s atmosphere is derived and maintained from the photosynthesis of green plants. In a very real sense, O2 is the signature of life on Earth (Sagan et al. 1993).

The century-old science of biogeochemistry recognizes that the influence of life is so pervasive that there is no pure science of geochemistry at the surface of Earth (Vernadsky 1998). Indeed, many of the Earth’s characteristics are only hospitable to life today because of the current and historic abundance of life on this planet (Reiners 1986). Granted some Earthly characteristics, such as its gravity, the seasons, and the radiation received from the Sun, are determined by the size and position of our planet in the solar system. But most other features, including liquid water, climate, and a nitrogen-rich atmosphere, are at least partially due to the presence of life. Life is the bio in biogeochemistry.

At present, there is ample evidence that our species, Homo sapiens, is leaving unusual imprints on Earth’s chemistry. The human combustion of fossil fuels is raising the concentration of carbon dioxide in our atmosphere to levels not seen in the past 20 million years (Pearson and Palmer 2000). Our release of an unusual class of industrial compounds known as chlorofluorocarbons has depleted the concentration of ozone in the upper atmosphere, where it protects the Earth’s surface from harmful levels of ultraviolet light (Rowland 1989). In our effort to feed 7 billion people, we produce vast quantities of nitrogen and phosphorus fertilizers, resulting in the runoff of nutrients that pollute surface and coastal waters (Chapter 12). As a result of coal combustion and other human activities, the concentrations of mercury in freshly caught fish are much higher than a century ago (Monteiro and Furness 1997), rendering many species unfit for regular human consumption. Certainly we are not the first species that has altered the chemical environment of planet Earth, but if our current behavior remains unchecked, it is well worth asking if we may jeopardize our own persistence.

Understanding the earth as a chemical system

Just as a laboratory chemist attempts to observe and understand the reactions in a closed test tube, biogeochemists try to understand the chemistry of nature, where the reactants are found in a complex mix of materials in solid, liquid, and gaseous phases. In most cases, biogeochemistry is a nightmare to a traditional laboratory chemist: the reactants are impure, their concentrations are low, and the temperature is variable. About all you can say about the Earth as a chemical system is that it is closed with respect to mass, save for a few meteors arriving and a few satellites leaving our planet. This closed chemical system is powered by the receipt of energy from the Sun, which has allowed the elaboration of life in many habitats (Falkowski et al. 2008).

Biogeochemists often build models for what controls Earth’s surface chemistry and how Earth’s chemistry may have changed through the ages. Unlike laboratory chemists, we have no replicate planets for experimentation, so our models must be tested and validated by inference. If our models suggest that the accumulation of organic materials in ocean sediments is associated with the deposition of gypsum (CaSO4·2H2O), we must dig down through the sedimentary layers to see if this correlation occurs in the geologic record (Garrels and Lerman 1981). Finding the correlation does not prove the model, but it adds a degree of validity to our understanding of how Earth works—its biogeochemistry. Models must be revised when observations are inconsistent with their predictions.

Earth’s conditions, such as the composition of the atmosphere, change only slowly from year to year, so biogeochemists often build steady-state models. As an example, in a steady-state model of the atmosphere, the inputs and losses of gases are balanced each year; the individual molecules in the atmosphere change, but the total content of each stays relatively constant. The assumption of a steady-state brings a degree of tidiness to our models of Earth’s chemistry, but we should always be cognizant of the potential for nonlinear and cyclic behavior in Earth’s characteristics. Indeed, some cycles, such as the daily rotation of the Earth around its axis and its annual rotation about the Sun, are now so obvious that it seems surprising that they were mysterious to philosophers and scientists throughout much of human history.

Steady-state models often are unable to incorporate the cyclic activities of the biosphere, which we define as the sum of all the live and dead materials on Earth.¹ During the summer, total plant photosynthesis in the Northern Hemisphere exceeds respiration by decomposers. This results in a temporary storage of carbon in plant tissues and a seasonal decrease in atmospheric CO2, which is lowest during August of each year in the Northern Hemisphere (Figure 1.1). The annual cycle is completed during the winter months, when atmospheric CO2 returns to higher levels, as decomposition continues when many plants are dormant or leafless. Certainly, it would be a mistake to model the activity of the biosphere by considering only the summertime conditions, but a steady-state model can ignore the annual cycle if it uses a particular time each year as a baseline condition to examine changes over decades.

Figure 1.1 Annual cycles of CO 2 and O 2 in the atmosphere. Changes in the concentration of O 2 are expressed relative to concentrations of nitrogen (N 2 ) in the same samples. Note that the peak of O 2 in the atmosphere corresponds to the minimum CO 2 in late summer, presumably due to the seasonal course of photosynthesis in the Northern Hemisphere.

Source: From Ralph Keeling, unpublished data used by permission.

Over a longer time frame, the size of the biosphere has decreased during glacial periods and increased during post-glacial recovery. Similarly, the storage of organic carbon increased strongly during the Carboniferous Period—about 300 million years ago, when most of the major deposits of coal were laid down. The unique conditions of the Carboniferous Period are poorly understood, but it is certainly possible that such conditions are part of a long-term cycle that might return again. Significantly, unless we recognize the existence and periodicity of cycles and nonlinear behavior and adjust our models accordingly we may err in our assumption of a steady state in Earth’s biogeochemistry.

All current observations of global change must be evaluated in the context of underlying cycles and potentially non-steady-state conditions in the Earth’s system. The current changes in atmospheric CO2 are best viewed in the context of cyclic changes seen during the last 800,000 years in a record obtained from the bubbles of air trapped in the Antarctic ice pack. These bubbles have been analyzed in a core taken near Vostok, Antarctica (Figure 1.2). During the entire 800,000-year period, the concentration of atmospheric CO2 appears to have oscillated between high values during warm periods and lower values during glacial intervals. Glacial cycles are linked to small variations in Earth’s orbit that alter the receipt of radiation from the Sun (Berger 1978; Harrington 1987). During the peak of the last glacial epoch (20,000 years ago), CO2 ranged from 180 to 200 ppm in the atmosphere. CO2 rose dramatically at the end of the last glacial (10,000 years ago) and was relatively stable at 280 ppm until the Industrial Revolution. The rapid increase in CO2 at the end of the last glacial epoch may have amplified the global warming that melted the continental ice sheets (Sowers and Bender 1995, Shakun et al. 2012).

Figure 1.2 An 800,000-year record of CO 2 and temperature, showing the minimum temperatures correspond to minimum CO 2 concentrations seen in cycles of ~120,000 periodicity, associated with Pleistocene glacial epochs.

Source: From Luthi et al. (2008)

When viewed in the context of this cycle, we can see that the recent increase in atmospheric CO2 to today’s value of about 400 ppm has occurred at an exceedingly rapid rate, which carries the planet into a range of concentrations never before experienced during the evolution of modern human social and economic systems, starting about 8000 years ago (Flückiger et al. 2002). If the past is an accurate predictor of the future, higher atmospheric CO2 will lead to global warming, but any observed changes in global climate must also be evaluated in the context of long-term cycles in climate with many possible causes (Crowley 2000; Stott et al. 2000).

The Earth has many feedbacks that buffer perturbations of its chemistry, so that steady-state models work well under many circumstances. For instance, Robert Berner and his coworkers at Yale University have elucidated the components of a carbonate–silicate cycle that stabilizes Earth’s climate and its atmospheric chemistry over geologic time (Berner and Lasaga 1989). The model is based on the interaction of carbon dioxide with Earth’s crust. Since CO2 in the atmosphere dissolves in rainwater to form carbonic acid (H2CO3), it reacts with the minerals exposed on land in the process known as rock weathering (Chapter 4). The products of rock weathering are carried by rivers to the sea (Figure 1.3).

Figure 1.3 The interaction between the carbonate and the silicate cycles at the surface of Earth. Long-term control of atmospheric CO 2 is achieved by dissolution of CO 2 in surface waters and its participation in the weathering of rocks. This carbon is carried to the sea as bicarbonate , and it is eventually buried as part of carbonate sediments in the oceanic crust. CO 2 is released back to the atmosphere when these rocks undergo metamorphism at high temperature and pressures deep in Earth.

Source: Modified from Kasting et al. (1988.

In the oceans, limestone (calcium carbonate) and organic matter are deposited in marine sediments, which in time are carried by subduction into Earth’s upper mantle. Here the sediments are metamorphosed; calcium and silicon are converted back into the minerals of silicate rock, and the carbon is returned to the atmosphere as CO2 in volcanic emissions. On Earth, the entire oceanic crust appears to circulate through this pathway in <200 million years (Muller et al. 2008). The presence of life on Earth does not speed the turning of this cycle, but it may increase the amount of material moving in the various pathways by increasing the rate of rock weathering on land and the rate of carbonate precipitation in the sea.

The carbonate–silicate model is a steady-state model, in the sense that it shows equal transfers of material along the flow-paths and no change in the mass of various compartments over time. In fact, such a model suggests a degree of self-regulation of the system, because any period of high CO2 emissions from volcanoes should lead to greater rates of rock weathering, removing CO2 from the atmosphere and restoring balance to the system. However, the assumption of a steady state may not be valid during transient periods of rapid change. For example, high rates of volcanic activity may have resulted in a temporary increase in atmospheric CO2 and a period of global warming during the Eocene, 40 million years ago (Owen and Rea 1985). And clearly, since the Industrial Revolution, humans have added more carbon dioxide to the atmosphere than the carbonate–silicate cycle or the ocean can absorb each year (Chapter 11).

Because the atmosphere is well mixed, changes in its composition are perhaps our best evidence of human alteration of Earth’s surface chemistry. Concern about global change is greatest when we see increases in atmospheric content of constituents such as carbon dioxide, methane (CH4), and nitrous oxide (N2O), for which we see little or no precedent in the geologic record. These gases are produced by organisms, so changes in their global abundance must reflect massive changes in the composition or activity of the biosphere.

Humans have also changed other aspects of Earth’s natural biogeochemistry. For example, when human activities increase the erosion of soil, we alter the natural rate of sediment delivery to the oceans and the deposition of sediments on the seafloor (Wilkinson and McElroy 2007, Syvitski et al. 2005). As in the case of atmospheric CO2, evidence for global changes in erosion induced by humans must be considered in the context of long-term oscillations in the rate of crustal exposure, weathering, and sedimentation due to changes in climate and sea level (Worsley and Davies 1979, Zhang et al. 2001).

Human extraction of fossil fuels and the mining of metal ores substantially enhance the rate at which these materials are available to the biosphere, relative to background rates dependent on geologic uplift and surface weathering (Bertine and Goldberg 1971). For example, the mining and industrial use of lead (Pb) has increased the transport of Pb in world rivers by about a factor of 10 (Martin and Meybeck 1979). Recent changes in the content of lead in coastal sediments appear directly related to fluctuations in the use of Pb by humans, especially in leaded gasoline (Trefry et al. 1985)—trends superimposed on underlying natural variations in the movements of Pb at Earth’s surface (Marteel et al. 2008, Pearson et al. 2010).

Recent estimates suggest that the global cycles of many metals have been significantly increased by human activities (Table 1.1). Some of these metals are released to the atmosphere and deposited in remote locations (Boutron et al. 1994). For example, the combustion of coal has raised the concentration of mercury (Hg) deposited in Greenland ice layers in the past 100 years (Weiss et al. 1971). Recognizing that the deposition of Hg in the Antarctic ice cap shows large variations over the past 34,000 years (Vandal et al. 1993), we must evaluate any recent increase in Hg deposition in the context of past cyclic changes in Hg transport through the atmosphere. Again, human-induced changes in the movement of materials through the atmosphere must be placed in the context of natural cycles in Earth system function (Nriagu 1989).

Table 1.1 Movement of Selected Elements through the Atmosphere

Source: From Lantzy and MacKenzie (1979). Used with permission.

aAll data are expressed in 10⁸ g/yr.

Scales of endeavor

The science of biogeochemistry spans a huge range of space and time, spanning most of the geologic epochs of Earth’s history (see inside back cover). Molecular biologists contribute their understanding of the chemical structure and spatial configuration of biochemical molecules, explaining why some biochemical reactions occur more readily than others (Newman and Banfield 2002). Increasingly, genomic sequencing allows biogeochemists to identify the microbes that are active in soils and sediments and what regulates their gene expression (Running et al. 2004). Physiologists measure variations in the activities of organisms, while ecosystem scientists measure the movement of materials and energy through well-defined units of the landscape.

Geologists study the chemical weathering of minerals in rocks and soils and document Earth’s past from sedimentary cores taken from lakes, oceans, and continental ice packs. Atmospheric scientists provide details of reactions between gases and the radiative properties of the planet. Meanwhile, remote sensing from aircraft and satellites allows biogeochemists to see the Earth at the largest scale, measuring global photosynthesis (Running et al. 2004) and following the movement of desert dusts around the planet (Uno et al. 2009). Indeed the skills needed by the modern biogeochemist are so broad that many students find their entrance to this new field bewildering. But the fun of being a biogeochemist stems from the challenge of integrating new science from diverse disciplines. And luckily, there are a few basic rules that guide the journey, as described in the next few subsections.

Thermodynamics

Two basic laws of physical chemistry, the laws of thermodynamics, tell us that energy can be converted from one form to another and that chemical reactions should proceed spontaneously to yield the lowest state of free energy, G, in the environment. The lowest free energy of a chemical reaction represents its equilibrium, and it is found in a mix of chemical species that show maximum bond strength and maximum disorder among the components. In the face of these basic laws, living systems create non-equilibrium conditions; life captures energy to counteract reactions that might happen spontaneously to maximize disorder.

Even the simplest cell is an ordered system; a membrane separates an inside from an outside, and the inside contains a mix of very specialized molecules. Biological molecules are collections of compounds with relatively weak bonds. For instance, to break the covalent bonds between two carbon atoms requires 83 kcal/mole, versus 192 kcal/mole for each of the double bonds between carbon and oxygen in CO2 (Davies 1972, Morowitz 1968). In living tissue most of the bonds between carbon (C), hydrogen (H), nitrogen (N), oxygen (O), phosphorus (P), and sulfur (S), the major biochemical elements, are reduced or electron-rich bonds that are relatively weak (Chapter 7). It is an apparent violation of the laws of thermodynamics that the weak bonds in the molecules of living organisms exist in the presence of a strong oxidizing agent in the form of O2 in the atmosphere. Thermodynamics would predict a spontaneous reaction between these components to produce CO2, H2O, and NO3−—molecules with much stronger bonds. In fact, after the death of an organism, this is exactly what happens! Living organisms must continuously process energy to counteract the basic laws of thermodynamics that would otherwise produce disordered systems with oxidized molecules and stronger bonds.

During photosynthesis, plants capture the energy in sunlight and convert the strong bonds between carbon and oxygen in CO2 to the weak, reduced biochemical bonds in organic materials. As heterotrophic organisms, herbivores eat plants to extract this energy by capitalizing on the natural tendency for electrons to flow from reduced bonds back to oxidizing substances, such as O2. Heterotrophs oxidize the carbon bonds in organic matter and convert the carbon back to CO2. A variety of other metabolic pathways have evolved using transformations among other compounds (Chapters 2 and 7), but in every case metabolic energy is obtained from the flow of electrons between compounds in oxidized or reduced states. Metabolism is possible because living systems can sequester high concentrations of oxidized and reduced substances from their environment. Without membranes to compartmentalize living cells, thermodynamics would predict a uniform mix, and energy transformations, such as respiration, would be impossible.

Free oxygen appeared in Earth’s surface environments sometime after the appearance of autotrophic, photosynthetic organisms (Chapter 2). Free O2 is one of the most oxidizing substances known, and the movement of electrons from reduced substances to O2 releases large amounts of free energy. Thus, large releases of free energy are found in aerobic metabolism, including the efficient metabolism of eukaryotic cells. The appearance of eukaryotic cells on Earth was not immediate; the fossil record suggests that they evolved nearly 1.5 billion years after the appearance of the simplest living cells (Knoll 2003). Presumably the evolution of eukaryotic cells was possible only after the accumulation of sufficient O2 in the environment to sustain aerobic metabolic systems. In turn, aerobic metabolism offered large amounts of energy that could allow the elaborate structure and activity of higher organisms. Here some humility is important: eukaryotic cells may perform biochemistry faster and more efficiently, but the full range of known biochemical transformations is found amongst the members of the prokaryotic kingdom.

Stoichiometry

A second organizing principle of biogeochemistry stems from the coupling of elements in the chemical structure of the molecules of which life is built—cellulose, protein, and the like. Redfield’s (1958) observation of consistent amounts of C, N, and P in phytoplankton biomass is now honored by a ratio that carries his name (Chapter 9). Reiners (1986) carried the concept of predictable stoichiometric ratios in living matter to much of the biosphere, allowing us to predict the movement of one element in an ecosystem by measurements of another. Sterner and Elser (2002) have presented stoichiometry as a major control on the structure and function of ecosystems. The growth of land plants is often determined from the nitrogen content of their leaves and the nitrogen availability in the soil (Chapter 6), whereas phosphorus availability explains much of the variation in algal productivity in lakes (Chapter 8). The population growth of some animals may be determined by sodium—an essential element that is found at a low concentration in potential food materials, relative to its concentration in body tissues.

Although the stoichiometry of biomass allows us to predict the concentration of elements in living matter, the expected ratio of elements in biomass is not absolute such as the ratio of C to N in a reagent bottle of alanine. For instance, a sample of phytoplankton will contain a mix of species that vary in individual N/P ratios, with the weighted average close to that postulated by Redfield (Klausmeier et al. 2004). And, of course, a large organism will contain a mix of metabolic compounds (largely protein) and structural components (e.g., wood or bone) that differ in elemental composition (Reiners 1986, Arrigo et al. 2005; Elser et al. 2010). In some sense, organisms are what they eat, but decomposers can adjust their metabolism (Manzoni et al. 2008) and enzymatic production (Stock et al. 1990) to feed on a wide range of substrates, even as they maintain a constant stoichiometry in their own biomass.

In some cases, trace elements control the cycle of major elements, such as nitrogen, by their role as activators and cofactors of enzymatic synthesis and activity. When nitrogen supplies are low, signal transduction by P activates the genes for N fixation in bacteria (Stock et al. 1990). The enzyme for nitrogen fixation, nitrogenase, contains iron (Fe) and molybdenum (Mo). Over large areas of the oceans, Falkowski et al. (1998) show that iron, delivered to the surface waters by the wind erosion of desert soils, controls marine production, which is often limited by N fixation. Similarly, when phosphorus supply is low, plants and microbes may produce alkaline phosphatase, containing zinc, to release P from dead materials (Shaked et al. 2006). Thus, the productivity of some ecosystems can be stimulated either by adding the limiting element itself or by adding a trace element that facilitates nutrient acquisition (Arrigo et al. 2005).

The elements of life are also coupled in metabolism, since organisms employ some elements in energy-yielding reactions, without incorporating them into biomass. Coupled biogeochemistry of the elements in metabolism stems from the flow of electrons in the oxidation/reduction reactions that power all of life (Morowitz 1968, Falkowski et al. 2008). Coupled metabolism is illustrated by a matrix, where each element in a column is reduced while the element in an intersecting row is oxidized (Figure 1.4). All of Earth’s metabolisms can be placed in the various cells of this matrix and in a few adjacent cells that would incorporate columns and rows for Fe and other trace metals. The matrix incorporates the range of metabolisms possible on Earth, should the right conditions exist (Bartlett 1986).

Figure 1.4 A matrix showing how cellular metabolisms couple oxidation and reduction reactions. The cells in the matrix are occupied by organisms or a consortium of organisms that reduce the element at the top of the column, while oxidizing an element at the beginning of the row.

Source: From Schlesinger et al.(2011).

Large-Scale Experiments

Biogeochemists frequently conduct large-scale experiments to assess the response of natural systems to human perturbation. Schindler (1974) added phosphorus to experimental lakes in Canada to show that it was the primary nutrient limiting algal growth in those ecosystems (Figure 1.5). Bormann et al. (1974) deforested an entire watershed to demonstrate the importance of vegetation in sequestering nutrients in ecosystems. Several experiments have exposed replicated plots of forests, grasslands, and desert ecosystems to high CO2 to simulate plant growth in the future environments on Earth (Chapter 5). And oceanographers have added Fe to large patches of the sea to ascertain whether it normally limits the growth of marine phytoplankton (Chapter 9). In many cases these large experiments and field campaigns are designed to test the predictions of models and to validate them.

Figure 1.5 An ecosystem-level experiment in which a lake was divided and one half (distant) fertilized with phosphorus, while the basin in the foreground acted as a control. The phosphorus-fertilized basin shows a bloom of nitrogen-fixing cyanobacteria.

Source: From Schindler (1974); www.sciencemag.org/content/184/4139/897.short. Used with permission.

Models

With sufficient empirical observations, biogeochemists can often build mathematical models for how ecosystems function. Equations can express what controls the movement of energy and materials through organisms or individual compartments of an ecosystem, such as the soil. The equations often incorporate the constraints of thermodynamics and stoichiometry. These models allow us to determine which processes control the productivity and biogeochemical cycling in ecosystems, and where our understanding is incomplete. Models that are able to reproduce past dynamics reliably allow us to explore the behavior of ecosystems in response to future perturbations that may lie outside the natural range of environmental variation.

Lovelock’s gaia

In a provocative book, Gaia, published in 1979, James Lovelock focused scientific attention on the chemical conditions of the present-day Earth, especially in the atmosphere, that are extremely unusual and in disequilibrium with respect to thermodynamics. The 21% atmospheric content of O2 is the most obvious result of living organisms, but other gases, including NH3 and CH4, are found at higher concentrations than one would expect in an O2-rich atmosphere (Chapter 3). This level of O2 in our atmosphere is maintained despite known reactions that should consume O2 in reaction with crustal minerals and organic carbon. Further, Lovelock suggested that the albedo (reflectivity) of Earth must be regulated by the biosphere, because the planet has shown relatively small changes in surface temperature despite large fluctuations in the Sun’s radiation during the history of life on Earth (Watson and Lovelock 1983).

Lovelock suggested that the conditions of our planet are so unusual that they could only be expected to result from activities of the biosphere. Indeed, Gaia suggests that the biosphere evolved to regulate conditions within a range favorable for the continued persistence of life on Earth. In Lovelock’s view, the planet functions as a kind of superorganism, providing planetary homeostasis. Reflecting the vigor and excitement of a new scientific field, other workers have strongly disagreed—not denying that biotic factors have strongly influenced the conditions on Earth, but not accepting the hypothesis of purposeful self-regulation of the planet (Lenton 1998).

Like all models, Gaia remains as a provocative hypothesis, but the rapid pace at which humans are changing the biosphere should alarm us all. Some ecologists see the potential for critical transitions in ecosystem function; points beyond which human impacts would not allow the system to rebound to its prior state, even if the impacts ceased (Scheffer et al. 2009). Others have attempted to quantify these thresholds, so that we may recognize them in time (Rockstrom et al. 2009). In all these endeavors, policy makers are desperate for biogeochemists to deliver a clear articulation of how the world works, the extent and impact of the human perturbation, and what to do about it.

Recommended Readings

Gorham, E. 1991. Biogeochemistry: Its origins and development. Biogeochemistry 13:199–239.

Kump, L.R., J.F. Kasting, and R.G. Crane. 2010. The Earth System, second ed. Prentice Hall.

Lovelock, J.E. 2000. The Ages of Gaia. Oxford University Press.

Sterner, R.W., and J.J. Elser. 2002. Ecological Stoichiometry. Princeton University Press.

Smil, V. 1997. The Cycles of Life.Scientific American Press.

Volk, T. 1998. Gaia’s Body. Springer/Copernicus.

Williams, G.R. 1996. The Molecular Biology of Gaia. Columbia University Press.

¹Some workers use the term biosphere to refer to the regions or volume of Earth that harbor life. We prefer the definition used here, so that the oceans, atmosphere, and surface crust can be recognized separately. Our definition of the biosphere recognizes that it has mass, but also functional properties derived from the species that are present.

Chapter 2

Origins

Outline

Introduction

Origins of the Elements

Origin of the Solar System and the Solid Earth

Origin of the Atmosphere and the Oceans

Origin of Life

Evolution of Metabolic Pathways

Photosynthesis: The Origin of Oxygen on Earth

Chemoautotrophy

Anaerobic Respiration

Comparative Planetary History: Earth, Mars, and Venus

Summary

Introduction

Six elements, H, C, N, O, P, and S, are the major constituents of living tissue and account for 95% of the mass of the biosphere. At least 25 other elements are known to be essential to at least one form of life, and it is possible that this list may grow slightly as we improve our understanding of the role of trace elements in biochemistry (Williams and Fraústo da Silva 1996).¹ In the periodic table (see inside front cover), nearly all the elements essential to life are found at atomic numbers lower than that of iodine at 53. Even though living organisms affect the distribution and abundance of some of the heavier elements, the biosphere is built from the light elements (Deevey 1970, Wackett et al. 2004). Ultimately, the environment in which life arose and the arena for biogeochemistry today was determined by the relative abundance of chemical elements in our galaxy and by the subsequent concentration and redistribution of those elements on Earth’s surface.

In this chapter we will examine models that astrophysicists suggest for the origin of the elements. Then we will examine models for the formation of the solar system and the planets. There is good evidence that the conditions on the surface of the Earth changed greatly during the first billion years or so after its formation—before life arose. Early differentiation of the Earth, the cooling of its surface, and the composition of the earliest oceans determined the arena for the origins of life. Later changes caused by the evolution and proliferation of life strongly determined the conditions on our planet today. In this chapter, we will consider the origin of the major metabolic pathways that characterize life and affect Earth’s biogeochemistry. The chapter ends with a discussion of the planetary evolution that has occurred on Earth compared to its near neighbors—Mars and Venus.

Origins of the Elements

Any model for the origin of the chemical elements must account for their relative abundance in the Universe. Estimates of the cosmic abundance of elements are made by examining the spectral emission from the stars in distant galaxies as well as the emission from our Sun (Ross and Aller 1976). Analyses of meteorites also provide important information on the composition of the solar system (Figure 2.1). Two points are obvious: (1) with three exceptions—lithium (Li), beryllium (Be), and boron (B)—the light elements, that is, those with an atomic number <30, are far more abundant than the heavy elements; (2) especially among the light elements, the even-numbered elements are more abundant than the odd-numbered elements of similar atomic weight.

Figure 2.1 The relative abundance of elements in the solar system, also known as the cosmic abundance, as a function of atomic number. Abundances are plotted logarithmically and scaled so that silicon (Si) = 1,000,000.

Source: From a drawing in Brownlee (1992) based on the data of Anders and Grevesse (1989).

A central theory of astrophysics is that the Universe began with a gigantic explosion, the Big Bang, about 13.7 billion years ago (Freedman and Madore 2010). The Big Bang initiated the fusion of hypothetical fundamental particles, known as quarks, to form protons (¹H) and neutrons, and it allowed the fusion of protons and neutrons to form some simple atomic nuclei (²H, ³He, ⁴He, and a small amount of ⁷Li). See Malaney and Fowler (1988), Pagel (1993), and Copi et al. (1995). After the Big Bang, the Universe began to expand outward, so there was a rapid decline in the temperatures and pressures that would be needed to produce heavier elements by fusion in interstellar space. Moreover, the elements with atomic masses of 5 and 8 are unstable, so no fusion of the abundant initial products of the Big Bang (i.e., ¹H and ⁴He) could yield an appreciable, persistent amount of a heavier element. Thus, the Big Bang can explain the origin of elements up to ⁷Li, but the origin of heavier elements had to await the formation of stars in the Universe—about 1 billion years later.

A model for the synthesis of heavier elements in stars was first proposed by Burbidge et al. (1957), who outlined a series of pathways that could occur in the interior of massive stars during their evolution (Fowler 1984, Wallerstein 1988, Trimble 1997). As a star ages, the abundance of hydrogen (H) in the core declines as it is converted to helium (He) by fusion. As the heat from nuclear fusion decreases, the star begins to collapse inward under its own gravity. This collapse increases the internal temperature and pressure until He begins to be converted via fusion reactions to form carbon (C) in a two-step reaction known as the triple-alpha process. First,

(2.1)

Then, while most ⁸Be decays spontaneously back to helium, the momentary existence of small amounts of ⁸Be under these conditions allows reaction with helium to produce carbon:

(2.2)

The main product of this so-called helium burning reaction is ¹²C, and the rate of this reaction determines the abundance of C in the Universe (Oberhummer et al. 2000). ¹⁶O is built by the addition of ⁴He to ¹²C, and nitrogen by the successive addition of protons to ¹²C. As the supply of helium begins to decline, a second phase of stellar collapse is followed by the initiation of a sequence of further fusion reactions in massive stars (Fowler 1984). First, fusion of two ¹²C forms ²⁴Mg (magnesium), some of which decays to ²⁰Ne (neon) by loss of an alpha (⁴He) particle. Subsequently, oxygen burning produces ³²S, which forms an appreciable amount of ²⁸Si (silicon) by loss of an alpha particle (Woosley 1986).

A variety of fusion reactions in massive stars are thought to be responsible for the synthesis—known as stellar nucleosynthesis—of the even-numbered elements up to iron (Fe) (Fowler 1984, Trimble 1997). (Smaller stars, like our Sun, do not go through all these reactions and burn out along the way, becoming white dwarfs.) These fusion reactions release energy and produce increasingly stable nuclei (Friedlander et al. 1964). However, to make a nucleus heavier than Fe requires energy, so when a star’s core is dominated by Fe, it can no longer burn. This leads to the catastrophic collapse and explosion of the star, which we recognize as a supernova. Heavier elements are apparently formed by the successive capture of neutrons by Fe, either deep in the interior of stable stars (s-process) or during the explosion of a supernova (r-process; Woosley and Phillips 1988, Burrows 2000, Cowan and Sneden 2006). A supernova casts all portions of the star into space as hot gases (Chevalier and Sarazin 1987).

This model explains a number of observations about the abundance of the chemical elements in the Universe. First, the abundance of elements declines logarithmically with increasing mass beyond hydrogen and helium, the original building blocks of the Universe. However, as the Universe ages, more and more of the hydrogen will be converted to heavier elements during the evolution of stars. Astrophysicists can recognize younger, second-generation stars, such as our Sun, that have formed from the remnants of previous supernovas because they contain a higher abundance of iron and heavier elements than older, first-generation stars, in which the initial hydrogen-burning reactions are still predominant (Penzias 1979). We should all be thankful for the fusion reactions in massive stars which have formed most of the chemical elements of life.

Second, because the first step in the formation of all the elements beyond lithium is the fusion of nuclei with an even number of atomic mass (e.g., ⁴He, ¹²C), the even-numbered light elements are relatively abundant in the cosmos. The odd-numbered light elements are formed by the addition of neutrons to nuclei in the interior of massive stars (s-process) and by the fission of heavier even-numbered nuclei. In most cases an odd-numbered nucleus is slightly less stable than its even-numbered neighbors, so we should expect odd-numbered nuclei to be less abundant. For example, phosphorus is formed in the reaction

(2.3)

Thus, phosphorus is much less abundant than the adjacent elements in the periodic table, Si (silicon) and S (sulfur) (Figure 2.1). It is interesting to speculate that the low cosmic abundance of P (phosphorus) formed by this and other reactions of nucleosynthesis may account for the fact that P is often in short supply for the biosphere on Earth today (Macia et al. 1997).

Finally, the low cosmic abundance of Li, Be, and B is due to the fact that the initial fusion reactions pass over nuclei with atomic masses of 5 to 8, forming ¹²C, as shown in Eqs. 2.1 and 2.2. Apparently, most Li, Be, and B are formed by spallation—the fission of heavier elements that are hit by cosmic rays in interstellar space (Olive and Schramm 1992, Reeves 1994, Chaussidon and Robert 1995).

This model for the origin and cosmic abundance of the elements offers some initial constraints for biogeochemistry. All things being equal, we might expect that the chemical environment in which life arose would approximate the cosmic abundance of elements. Thus, the evolution of biochemical molecules might be expected to capitalize on the light elements that were abundant in the primordial environment. It is then of no great surprise that no element heavier than Fe is more than a trace constituent in living tissue and that among the light elements, no Li or Be, and only traces of B, are essential components of biochemistry (Wackett et al. 2004). The composition of life is remarkably similar to the composition of the Universe; as put by Fowler (1984), we are all a little bit of stardust.

Origin of the solar system and the solid earth

The Milky Way galaxy is about 12.5 billion years old (Dauphas 2005), indicating that the first stars and galaxies had formed within a billion years after the Big Bang (Cayrel et al. 2001). By comparison, as a second-generation star, our Sun appears to be only about 4.57 billion years old (Baker et al. 2005, Bonanno et al. 2002, Bouvier and Wadhawa 2010). Current models for the origin of the solar system suggest that the Sun and its planets formed from a cloud of interstellar gas and dust, possibly including the remnants of a supernova (Chevalier and Sarazin 1987). This cloud of material would have the composition of the cosmic mix of elements (Figure 2.1). As the Sun and the planets began to condense, each developed a gravitational field that helped capture materials that added to its initial mass. The mass concentrated in the Sun apparently allowed condensation to pressures that reinitiated the fusion of hydrogen to helium.

The planets of our solar system appear to have formed from the coalescing of dust to form small bodies, known as planetesimals, within the primitive solar cloud (Beckwith and Sargent 1996, Baker et al. 2005). Collisions among the planetesimals would have formed the planets. The process is likely to have been fairly rapid. Several lines of evidence show planetesimals forming during the first million years of the solar system (Srinivasan et al. 1999, Yin et al. 2002, Alexander et al. 2001), and most stars appear to lose their disk of gases and dust within 400 million years after their formation (Habing et al. 1999). Recent observations suggest that a similar process is now occurring around another star in our galaxy, ß Pictoris (Lagage and Pantin 1994, Lagrange et al. 2010), and Earth-size and larger planets have been detected around numerous other stars in our galaxy (Gaidos et al. 2007, Borucki et al. 2010, Lissauer et al. 2011).

Overall, the original solar nebula is likely to have been composed of about 98% gaseous elements (H, He, and noble gases), 1.5% icy solids (H2O, NH3, and CH4), and 0.5% rocky solid materials, but the composition of each planet was determined by its position relative to the Sun and the rate at which the planet grew (McSween 1989). The inner planets (Mercury, Venus, Earth, and Mars) seemed to have formed in an area where the solar nebula was very hot, perhaps at a temperature close to 1200 K (Boss 1988). Venus, Earth, and Mars are all depleted in light elements compared to the cosmic abundances, and they are dominated by silicate minerals that condense at high temperatures and contain large amounts of FeO (McSween 1989). The mean density of Earth is about 5.5 g/cm³. The high density of the inner planets contrasts with the lower average density of the larger, outer planets, known as gas giants, which captured a greater fraction of lighter constituents from the initial solar cloud (Table 2.1). Jupiter contains much hydrogen and helium. The average density of Jupiter is 1.3 g/cm³, and its overall composition does not appear too different from the solar abundance of elements (Lunine 1989, Niemann et al. 1996). Some astronomers have pointed out that the hydrogen-rich atmosphere on Jupiter is similar to the composition of brown dwarfs—stars that never ignited (Kulkarni 1997).

Table 2.1 Characteristics of the Planets

Source: From Broecker (1985, p. 73). Published by Lamont Dougherty Laboratory, Columbia University. Used with permission.

aThe mass of the Sun is 1.99 × 10³³ gm, 1000× the mass of Jupiter.

bDensity a planet would have in the absence of gravitational squeezing.

From the initial solar cloud of elements, the chemical composition of the Earth is a selective mix, peculiar to the orbit of the incipient planet. The majority of the mass of the Earth seems likely to have accreted by about 4.5 bya—within about 100 million years of the origin of the solar system (Allègre et al. 1995, Kunz et al. 1998, Yin et al. 2002, Touboul et al. 2007, Jackson et al. 2010). Several theories account for the origin and differentiation of Earth. One suggests that Earth may have grown by homogeneous accretion; that is, throughout its early history, Earth may have captured planetesimals that were relatively similar in composition (Stevenson 1983, 2008).

Kinetic energy generated during the collision of these planetesimals (Wetherill 1985), as well as the heat generated from radioactive decay in its interior (Hanks and Anderson 1969), would heat the primitive Earth to the melting point of iron, nickel, and other metals, forming a magma ocean. These heavy elements were smelted from the materials arriving from space and sank to the interior of the Earth to form the core (Agee 1990, Newsom and Sims 1991, Wood et al. 2006).

As Earth cooled, lighter minerals progressively solidified to form a mantle dominated by perovskite (MgSiO3), with some complement of olivine (FeMgSiO4), and a crust dominated by aluminosilicate minerals of lower density and the approximate composition of feldspar (Chapter 4). Thus, despite the abundance of iron in the cosmos and in the Earth as a whole, the crust of the Earth is largely composed of Si, Al, and O (Figure 2.2). The aluminosilicate rocks of the crust float on the heavier semifluid rocks of the mantle (Figure 2.3; Bowring and Housh 1995).

Figure 2.2 Relative abundance of elements by weight in the whole Earth (a) and Earth’s crust (b).

Source: From Earth (fourth ed.) by Frank Press and Raymond Siever. Copyright 1986 by W.H. Freeman and Company. Used with permission.

Figure 2.3 A geologic profile of the Earth’s surface. On land the crust is dominated by granitic rocks, largely composed of Si and Al ( Chapter 4 ). The oceanic crust is dominated by basaltic rocks with a large proportion of Si and Mg. Both granite and basalt have a lower density than the upper mantle, which contains ultrabasic rocks with the approximate composition of olivine (FeMgSiO 4 ).

Source: From Howard and Mitchell (1985).

An alternative theory for the origin of Earth suggests that the characteristics of planetesimals and other materials contributing to the growth of the planet were not uniform through time. Theories of heterogeneous accretion suggest that materials in the Earth’s mantle arrived later than those of the core (Harper and Jacobsen 1996, Schönbächler et al. 2010), and that a late veneer delivered by a class of meteors known as carbonaceous chondrites was responsible for most of the light elements and volatiles on Earth (Anders and Owen 1977, Wetherill 1994, Javoy 1997, Kramers 2003). The two accretion theories are not mutually exclusive; it is possible that a large fraction of the Earth mass was delivered by homogeneous accretion, followed by a late veneer of chondritic materials (Willbold et al. 2011).

It is likely that during its late accretion, Earth was impacted by a large body—known as Theia—which knocked a portion of the incipient planet into an orbit about it, forming the Moon (Lee et al. 1997). The Moon’s age is estimated at 4.527 billion years (Kleine et al. 2005). Earth’s early history was probably dominated by frequent large impacts, but based on the age distribution of craters on the Moon, it is postulated that most of the large impacts occurred before 2.0 bya (Neukum 1977, Cohen et al. 2000, Bottke et al. 2012). The present-day receipt of extraterrestrial materials (8 to 38 × 10⁹ g/yr; Taylor et al. 1998, Love and Brownlee 1993, Cziczo et al. 2001) is much too low to account for Earth’s mass (6 × 10²⁷ g), even if it has continued for all of Earth’s history.

Consistent with either theory are several lines of evidence that the primitive Earth was devoid of an atmosphere derived from the solar nebula—that is, a primary atmosphere. During its early history, the gravitational field on the small, accreting Earth would have been too weak to retain gaseous elements, and the incoming planetesimals were likely to have been too small and too hot to carry an envelope of volatiles. The impact of Theia is also likely to have blown away any volatiles that had accumulated in Earth’s atmosphere by that time. Today, volcanic emissions of some inert (noble) gases, such as ³He, ²⁰Ne (neon), and ³⁶Ar (argon), which are derived from the solar nebula, result from continuing degassing of primary volatiles that must have been delivered to the primitive Earth trapped in pockets (fluid inclusions) in accreting chondrites (Lupton and Craig 1981, Burnard et al. 1997, Jackson et al. 2010). Otherwise, the Earth’s atmosphere appears to be of secondary origin.

If a significant fraction of today’s atmosphere were derived from the original solar cloud, we might expect that its gases would exist in proportion to their solar abundances (refer to Figure 2.1). Here, ²⁰Ne is of particular interest because it is not produced by any known radioactive decay, it is too heavy to escape from Earth’s gravity, and as an inert gas it is not likely to have been consumed in any reaction with crustal minerals (Walker 1977).² Thus, the present-day abundance of ²⁰Ne in the atmosphere is likely to represent its primary abundance—that derived from the solar nebula. Assuming that other solar gases were delivered to the Earth in a similar manner, we can calculate the total mass of the primary atmosphere by multiplying the mass of ²⁰Ne in today’s atmosphere by the ratio of each of the other gases to ²⁰Ne in the solar abundance. For example, the solar ratio of nitrogen to neon is 0.91 (Figure 2.1). If the present-day atmospheric mass of neon, 6.5 × l0¹⁶ g, is all from primary sources, then (0.91) × (6.5 × 10¹⁶ g) should be the mass of nitrogen that is also of primary origin. The product, 5.9

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This is a very thorough, broad, and detailed synthesis of the topic.