Wallace J.M., Hobbs P.V. Atmospheric Science. An Introductory Survey

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44 The Earth System

that settle to the ocean floor are converted into lime-

stone (CaCO

) rocks. This inorganic carbon reservoir

of the Earth’s crust is the largest of the carbon reser-

voirs in the Earth system.

2.3.3 Carbon in the Oceans

The carbon in the oceanic reservoir exists in three

forms: (1) dissolved CO

or H

, also known as

carbonic acid, (2) carbonate ions paired

with Ca

2

and Mg

2

and other metallic cations,

and (3) bicarbonate ions. The third

form is by far the largest of the oceanic carbon

reservoirs (Table 2.3). The dissolved carbon dioxide

concentration equilibrates with atmospheric con-

centrations through the reaction

(2.7)

An increase in the atmospheric concentration of CO

thus tends to raise the equilibrium concentration of

dissolved CO

. Carbonic acid, in turn, dissociates to

form bicarbonate ions and hydrogen ions

(2.8)

thereby causing the water to become more acidic.

The increasing concentration of H



ions shifts the

equilibrium between carbonate and bicarbonate ions

(2.9)

toward the left. The net effect, obtained by adding

(2.7), (2.8), and the reverse of (2.9) is

(2.10)

which incorporates the added carbon into the bicar-

bonate reservoir without any net increase in the

acidity of the ocean. The ability of the ocean to take

up and buffer CO

in this manner is limited by the

availability of ions in the carbonate reservoir.

Marine organisms incorporate bicarbonate ions

into their shells and skeletons through the reaction

(2.11)

A fraction of the calcium carbonate created in

Eq. (2.11) settles on the sea floor and forms lime-

stone deposits, while the remainder dissolves through

the reverse reaction.

2

 2HCO



: CaCO

 H

 CO

2

 H

O M 2HCO



HCO





 CO

2

M H



 HCO



 H

O M H

(HCO



)

(CO

2

)

(2.12)

Limestone deposits tend to be concentrated in con-

tinental shelves beneath shallow tropical seas that

provide an environment hospitable to coral. At these

levels in the ocean, the acidity of the water is low

enough that shells and skeletons deposited on the

ocean floor do not dissolve.

The Ca

2

ions that marine organisms incorporate

into their shells enter the ocean by way of the

weathering of rocks in rain water that is carried to

the oceans in rivers. Some of these ions are derived

from the weathering of calcium-silicate rocks in the

reaction

(2.13)

The net effect of Eqs. (2.11) and (2.13), in combina-

tion with Eq. (2.7), is

(2.14)

which has the effect of taking up CO

from the

atmospheric and oceanic reservoirs and incorporating

it into the much larger reservoir of inorganic carbon

sedimentary rocks in the Earth’s crust.

From a climate perspective, the chemical reactions

(2.7)–(2.14) are virtually instantaneous. In contrast,

the timescale over which the oceanic reservoirs

adjust to changes in atmospheric CO

is governed by

the ventilation time for the deeper layers of the

ocean, which is on the order of centuries. Calcium

carbonate formation is limited by the availability of

calcium ions, which is determined by the rate of

weathering of calcium silicate rocks, as described in

the following subsection.

2.3.4 Carbon in the Earth’s Crust

The organic and inorganic carbon reservoirs in the

Earth’s crust are both very large, the exchange rates

in and out of them (apart from the burning of fossil

fuels) are very slow, and residence times are on

the order of many millions of years. Carbon enters

both of these reservoirs by way of the biosphere,

as described earlier. Most deposits of the organic

carbon in natural gas, oil, coal, and shales and

other sedimentary rocks were formed in anoxic

ocean basins. The even larger inorganic carbon

reservoir, consisting mostly of calcium carbonate

CaSiO

 CO

: CaCO

 SiO

 H

CaSiO

 H

: Ca

2

 2HCO



CaCO

 H

: Ca

2

 2HCO



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2.4 Oxygen in the Earth System 45

(CaCO

), is almost exclusively a product of the

marine biosphere.

Weathering exposes organic carbon in sedimentary

rock to the atmosphere, allowing it to be oxidized,

thereby completing the loop in what is sometimes

referred to as the long term inorganic carbon cycle.

Currently the burning of fossil fuels is returning as

much carbon to the atmosphere in a single year as

weathering would return in hundreds of thousands

of years! The mass of carbon that exists in a form

concentrated enough to be classified as “fossil fuels”

represents only a small fraction of the organic carbon

stored in the Earth’s crust, but it is nearly an order of

magnitude larger than the mass of carbon currently

residing in the atmosphere.

On timescales of tens to hundreds of millions of

years, plate tectonics and volcanism play an essential

role in renewing atmospheric CO

. This “inorganic

carbon cycle,” summarized in Fig. 2.25, involves sub-

duction, metamorphism, and weathering. Limestone

sediments on the sea floor are subducted into the

Earth’s mantle along plate boundaries where conti-

nental plates are overriding denser oceanic plates. At

the high temperatures within the mantle, limestone is

transformed into metamorphic rocks by the reaction

(2.15)

The CO

released in this reaction eventually returns

to the atmosphere by way of volcanic eruptions. The

metamorphic rocks containing calcium in chemical

combination with silicate are recycled in the form of

newly formed crust that emerges in the mid-ocean

ridges. The metamorphism reaction (2.15), in combi-

nation with weathering, and the carbonate formation

reaction (2.14) form a closed loop in which carbon

atoms cycle back and forth between the atmospheric

CaCO

 SiO

: CaSiO

 CO

reservoir and the inorganic carbon reservoir in

the Earth’s crust on a timescale of tens to hundreds

of millions of years.

At times when the rate at which CO

is injected

into the atmosphere by volcanic eruptions exceeds

the rate at which calcium ions are made available by

weathering, atmospheric CO

concentrations increase

and vice versa. The injection rate is determined by

rate of metamorphism of carbonate rocks, which, in

turn, depends on the rate of plate movement along

convergent boundaries where subduction is occur-

ring. The rate of weathering, however, is proportional

to the rate of cycling of water in the atmospheric

branch of the hydrologic cycle, which increases with

increasing temperature. The fact that weathering

involves the chemical reaction (2.13) makes the

temperature dependence even stronger. Hence, high

ambient temperatures and slow plate movements

are conducive to a draw-down of atmospheric CO

and vice versa. The changes in atmospheric CO

response to imbalances between (2.14) and (2.15) on

timescales of tens of millions of years are believed to

have been quite substantial.

2.4 Oxygen in the Earth System

Earth is unique among the planets of the solar sys-

tem with respect to the abundance of atmospheric

oxygen O

and the presence of an ozone (O

) layer.

Atmospheric oxygen accounts for only a very small

fraction of the “free” oxygen (i.e., oxygen not bound

to hydrogen atoms in water molecules) in the Earth

system. Much larger quantities of free oxygen are

present in the form of oxidized minerals in sediments

and in the crust and upper mantle. The current level

of oxidation of the Earth system as a whole is much

higher than it was at the time when the planets first

formed.

The formation of the Earth’s molten metallic iron

had the effect of enriching oxygen concentrations in

the mantle, the source of volcanic emissions. Yet geo-

logical evidence suggests that oxygen was only a trace

atmospheric constituent early in the Earth’s history.

Iron in sedimentary rock formations that date back

more than 2.2 billion years is almost exclusively in the

partially oxidized ferrous (FeO) form. Had substantial

amounts of oxygen been present in the atmosphere

and oceans at the time when these sediments formed,

the iron in them would have been fully oxidized to

ferric oxide (Fe

). To account for the large concen-

trations of Fe

that currently reside in the Earth’s

Carbon Calcium Silicon

metamorphic

rocks

limestone

ions

quartz

oceans

atmosphere

Fig. 2.25 Schematic of the long-term inorganic carbon

cycle, also referred to as the carbonate–silicate cycle. The

symbol S denotes sedimentation, M denotes metamorphosis,

and W denotes weathering.

P732951-Ch02.qxd 9/12/05 7:40 PM Page 45

46 The Earth System

crust and upper mantle, it is necessary to invoke the

existence of a source of free oxygen that came into

play later in the history of the Earth system. Free

oxygen is also required to account for the formation of

carbonates in the Earth’s crust by marine organisms,

as illustrated in the following exercise.

Exercise 2.5 Estimate the mass of oxygen required

to form the carbonate deposits in the Earth’s crust.

Solution: From Table 2.3, the mass of carbon

in the carbonate reservoir in the Earth’s crust is

80,000 kg m

2

. Combining (2.7), (2.8), (2.9), and

(2.11), we obtain

The two hydrogen ions generated as a by-product of

this reaction combine with a single oxygen ion to

form water. Hence, the final result is

(2.16)

from which it is clear that one “free” oxygen atom is

paired with each carbon atom that resides in the

Earth’s crust in the form of carbonates. Hence, the mass

of free oxygen required to account for the carbonates is

or 100,000 kg m

2

, which is the basis for the estimate

in Table 2.4. ■

 80,000 kg m

2

 (16



12)

80,000 kg m

2



molecular weight of oxygen

molecular weight of carbon

 Ca

2

 O

2

: CaCO

O  CO

 Ca

2

: CaCO

 2H



2.4.1 Sources of Free Oxygen

The photosynthesis reaction (2.5) does not affect the

overall state of oxidation of the Earth system because

for every liberated O

molecule, an atom of organic

carbon is stored in sedimentary rocks in the Earth’s

crust. That carbon storage can be used to estimate

the mass of carbon liberated by photosynthesis, as

illustrated in the following exercise.

Exercise 2.6 On the basis of data in Table 2.3, esti-

mate the oxygen liberated by photosynthesis over

the lifetime of the Earth.

Solution: From Table 2.3, the mass of organic

carbon in sedimentary rocks in the Earth’s crust is

20,000 kg m

2

. From (2.5) it is evident that the burial

of each carbon atom in the Earth’s crust has liber-

ated one O

molecule. Hence, the mass of oxygen

liberated by photosynthesis is

or 50,000 kg m

2

. ■

Until quite recently, photosynthesis was believed to

have been the major source of free oxygen on Earth.

However, it is evident from Exercise 2.6 that a much

larger source is needed to account for the current

degree of oxidation of the Earth system.The only other

possible candidate is redox reactions that ultimately

lead to the escape of the hydrogen atoms to space, rais-

ing the state of oxidation of the Earth system. The rate

of escape of hydrogen atoms is proportional to their

number concentration in the upper atmosphere. The

principal source of free hydrogen is a four-step process.

1. Subduction or deep burial of rocks containing

hydrated minerals (i.e., minerals with chemically

bound H

O molecules or interstitial H

molecules in their lattices).

2. Breakdown of these hydrated minerals at

the high temperatures in the Earth’s mantle,

releasing the water in the form of steam.

3. The reaction

(2.17)

in which the steam oxidizes ferrous oxide in the

crust or mantle to ferric oxide.

4. The ejection of the hydrogen in volcanism or

metamorphism.

2FeO  H

O : Fe

 H

20,000 kg m

2



molecular weight of O

molecular weight of C

Table 2.4 Mass of oxygen (in units of 10

kg m

2

) averaged

over the surface of the Earth required to raise the level of

oxidation of the various reservoirs in the Earth system to

their present state, starting from their state at the time of

formation of the Earth

Reservoir Mass

Atmospheric O

2.353

Oceans and sediments 31

Crust Fe

3

100

Crust CO

100

Crust (other) 100

Mantle Fe

3

100

The estimates for the crust and mantle represent conservative lower limits.

[Based on data adapted with permission from Catling, D. C., K. J. Zahnle and

C. P. McKay, “Biogenic Methane, Hydrogen Escape, and the Irreversible

P732951-Ch02.qxd 9/12/05 7:40 PM Page 46

2.4 Oxygen in the Earth System 47

The escape of hydrogen from the Earth’s atmos-

phere is detectable (Fig. 2.26), but the rate at which it

is occurring is far too small to account for the degree

to which the minerals in the Earth’s crust and mantle

have become oxidized over the lifetime of the Earth.

The rate of escape is slow because the two gases that

supply hydrogen atoms to the upper atmosphere

(i.e., CH

and H

O) are present only in concen-

trations of a few parts per million by volume. Air

entering the upper atmosphere from below loses

most of its water vapor as it passes through the cold

equatorial tropopause (see Sections 1.3.3 and 3.5).

There are indications that atmospheric methane con-

centrations may have been much higher in the past,

as discussed in the next section.

Photosynthesis and the escape of hydrogen oper-

ate independently to liberate oxygen. The crust of a

lifeless planet could become highly oxidized due

solely to the action of plate tectonics. Conversely,

photosynthesis could occur on a planet on which the

minerals in the crust and the gases in the atmosphere

were in a highly reduced state. However, of the two

mechanisms, only photosynthesis is capable of pro-

ducing atmospheric oxygen, and only the escape of

hydrogen from the Earth system is capable of liber-

ating oxygen in the quantities required to account for

the present degree of oxidation of the minerals in the

crust and mantle.

Fig. 2.26 The corona around the Earth in this image is due to

the scattering of solar radiation by hydrogen atoms that are

escaping from the atmosphere. [Image from NASADynamics

Explorer/Spin-Scan Auroral Imaging. Courtesy of David Catling.]

Isotopes of a given element are atoms with differ-

ent numbers of neutrons in their nuclei. Unstable

isotopes, such as

C, which spontaneously change

form by radioactive decay with a known “half-

life,” are used for dating ice, tree and sediment

cores, fossils, and rock samples. By comparing the

abundance of the isotope in the sample with the

current atmospheric abundance, it is possible to

infer how long the sample has been out of contact

with the atmosphere (provided, of course, that the

atmospheric abundance has not changed since the

sample was deposited). Relative abundances of

stable isotopes, such as

C, vary in accordance

with local (and in some cases regional or even

global) environmental conditions that prevailed at

the time when the sample was deposited. Some of

the more widely used isotopes are the following.

• The relative abundance of deuterium (

or D)

in the snow samples recovered from

ice cores depends on (and hence can be used

as a proxy for) the temperature of the surface

from which the water vapor that condensed to

form the snow was evaporated.The greater

the relative abundance of HDO in the core

sample, the warmer that evaporating surface

2.1 Isotope Abundances: Proxies for Climate Data

The relative abundance of D is given by

where R is a reference value. Positive (negative) values of



D are indicative of an enrichment (depletion) of D relative to the reference

value, expressed in parts per thousand (ooo).The same formalism applies to the isotopes in subsequent bullets.



D(o/oo) 



H  R

Continued on next page

P732951-Ch02.qxd 9/12/05 7:40 PM Page 47

2.5 A Brief History of Climate

and the Earth System

This section describes evolution of the Earth system on

a logarithmically telescoping time line, as depicted in

the bottom panel of Fig. 2.27, with subsections focusing

on (1) the lifetime of the Earth, (2) the past 100 million

years, (3) the past million years, and (4) the past 20,000

years. If the life (to date) of a 20-year-old student were

viewed on an analogous and proportional time line, the

respective subsections would focus on hisher entire

20-year life span, the past 6 months, the past 2 days, and

the past hour.

2.5.1 Formation and Evolution

of the Earth System

The sun and the planets are believed to have formed

4.5 billion years ago from the gravitational collapse of

a cold cloud of interstellar gas and dust.

The absence

of the noble gases neon, xenon, and krypton

in the

48 The Earth System

must have been at the time the sample was

deposited.

• Oxygen-18 (

O) abundance in marine sediment

cores containing carbonates reflects the temper-

ature of the water in the euphotic zone where

the carbonate formation took place.The lower

the temperature of the water, the greater the

relative abundance of

C that was incorporated

into the shells and skeletons of the marine

organisms from which the carbonates formed.

• Worldwide

O abundances also depend on

the volume of the continental ice sheets.

evaporates more readily than

O, so a

disproportionately large fraction of

O tends

to be incorporated into the snow that falls on,

and becomes incorporated into, the ice sheets.

When the ice sheets grow, ocean waters

throughout the world are enriched in

Hence

O abundance in marine sediment

cores and ice cores can be used as a proxy for

ice volume.

• Carbon-13 (

C) abundances in deposits

of organic carbon reflect the ambient

concentrations at the time that

photosynthesis occurred. Plants prefer the

lighter isotope

C, and the higher the

ambient CO

level, the more strongly they

exert this preference. Hence, low relative

abundances of

C in organic carbon

deposits are indicative of high ambient

concentrations, and vice versa.

• Carbon-13 is also an indicator of the sources

and sinks of atmospheric CO

. Emissions

from the decay of plants, forest fires and

agricultural burning, and the consumption of

fossil fuels tend to be low in

C, whereas CO

outgassed from the oceans has the same

abundance as atmospheric CO

. In a similar

manner, the presence of a biospheric CO

sink should tend to raise atmospheric

levels, whereas the presence of an oceanic

sink should not.

2.1 Continued

Years before present × 10

lifetime of Earth

Hadean Epoch

origin of life

Years before present

lifetime of Earth

rise of oxygen

Pangaea

K-T boundary

plates collide, Himalayas rise

Panama closes gap

begin Quaternary Epoch

previous interglacial

last glacial maximum

Younger Dryas

“climatic optimum”

Medieval Warm Period

Little Ice Age

rise of oxygen

Pangaea

K-T boundary

Quaternary Epoch

Fig. 2.27 Time line for the history of the Earth system on a

linear scale (top) and on a logarithmic scale (bottom).

The age of the of the Earth is inferred from a comparison of the ratios of radiogenic (formed by decay of uranium) and nonradiogenic

isotopes of lead in meteorites and rocks of various ages.

Argon in the Earth’s atmosphere is a product of the radioactive decay of the radioactive isotope

K in the crust.

P732951-Ch02.qxd 9/12/05 7:40 PM Page 48

2.5 A Brief History of Climate and the Earth System 49

atmospheres of the Earth and the other planets, rela-

tive to their cosmic abundance, is evidence that the

planets formed from the coalescence of the dust into

chunks of solid materials called planetesimals that

were drawn together by gravitation. Present within the

condensing cloud were volatile compounds (i.e., water,

methane, ammonia, and other substances with low

boiling points), mainly in the form of ices. When the

sun formed, the inner part of the cloud should have

warmed, driving out most of the volatiles: hence the

relatively low concentrations of these substances in

the atmospheres of the inner planets.

During the first 700–800 million years of its history,

referred to by geologists as the Hadean Epoch, Earth

was still under continual bombardment by smaller

planetesimals. The heating and degassing resulting

from impacts of these collisions should have liberated

water vapor and other volatile substances, forming a

primordial atmosphere and the oceans. The energy

released by the larger objects may well have been suf-

ficient to entirely vaporize the oceans from time to

time. The formation of the moon has been attributed

to one of these impacts. The bombardment gradually

subsided and, by 3.8 billion years ago, conditions on

Earth had become stable enough to allow early micro-

bial life forms to develop in the oceans. Cataclysmic

collisions still occur occasionally, as evidenced by the

K–T meteorite impact

that took place only 65 million

years ago. Earth is still being bombarded by vast num-

bers of much smaller objects, as illustrated in Fig. 2.28,

but their impact on the Earth system is minimal.

The emission of radiation from stars increases

gradually over their lifetime due to the increasing

rate of fusion within their cores, which contract and

heat up as progressively more of the hydrogen is

fused into helium. The luminosity of the sun is

believed to have increased by 30% over the lifetime

of the solar system. Geological evidence indicates

that, with the exception of a few relatively brief

intervals, the oceans have been largely free of ice

throughout the Earth’s history. That the Earth’s

surface was not perpetually frozen during its early

history, when the sun was relatively faint, suggests

that its early atmosphere must have contained sub-

stantially higher concentrations of greenhouse gases

at that time than it does now.

Over the lifetime of the Earth, its atmosphere has

continuously been recycled and renewed by volcanism

and plate tectonics.The makeup of present-day volcanic

emissions is 80–90% steam, 6–12% CO

,1–2% SO

and traces of H

,CO,H

S, CH

, and N

. The relative

concentrations of the reduced gases H

, CO, and CH

could have been much higher earlier in the Earth’s his-

tory when the mantle was less oxidized than it is today.

An important milestone in the evolution of the

Earth system was the rise of atmospheric oxygen.

Cyano-(blue-green) bacteria capable of liberating

oxygen are believed to have been present in the oceans

for at least 3.0 and perhaps as long as 3.8 billion years,

yet geological evidence indicates that oxygen did not

begin to accumulate in the atmosphere until about

2.4–2.2 billion years ago. Oxygen liberated by photo-

synthesis early in the Earth’s history would have been

consumed quickly by H

and other reduced gases

emanating from the Earth’s crust, formed in reactions

such as (2.17) or by the oxidation of minerals exposed

to the atmosphere by weathering. Only after the

The distinctive marker of this event is the distinctive iridium-enriched layer produced by the explosion that attended the impact. This

layer, which is evident in sediments worldwide, occurs at the boundary between Cretaceous and Tertiary sediments in the Earth’s crust (hence

the name K–T). It has been hypothesized that this event was responsible for the extinction of many species of life forms, including dinosaurs.

–12

–10

–8

–6

(cm) (m) (km)

–4

–2

Years between impacts

Impactors on the space shuttle surface

(30 µ s)

(

µ m)

(30 s)

(1 year)

(10,000 years)

Shooting stars

Meteorites

Arizona crater

Sudbury, Ontario,

impact feature

Diameter of striking object (m)

Fig. 2.28 Frequency of occurrence of collisions of objects

with the Earth as a function of the size of the objects.

[Reprinted with permission from L. W. Alvarez, “Mass extinc-

tions caused by large bolide impacts,” Physics Today, 40, p. 27.

P732951-Ch02.qxd 9/12/05 7:40 PM Page 49

minerals in the crust became more highly oxidized

due to the gradual escape of hydrogen from the Earth

system and reactions such as (2.17) slowed down,

would it have it possible for oxygen liberated by

photosynthesis to begin to accumulate in the atmos-

phere. In this sense, atmospheric O

can be viewed as

“surplus” oxygen in the Earth system.

An array of geological evidence indicates that the

rise of atmospheric oxygen, once it began, was quite

rapid, with concentrations rising from less than 0.01%

of the present concentration 2.4 billion years ago to at

least 1–3% of the present concentration 1.9 billion

years ago. Concurrent with the rise of oxygen came

the formation of ozone layer. Photochemical models

based on the equations in Section 5.7.1 indicate that

atmospheric oxygen concentrations of even a few per-

cent of today’s values should have been capable of

supporting an ozone layer thick enough to protect life

on Earth from the harmful effects of solar ultraviolet

radiation.

The conditions that existed on Earth prior to the

rise of oxygen have been the subject of considerable

speculation. Given today’s rates of production of

methane, it is estimated that, in the absence of atmos-

pheric oxygen, methane concentrations could have

been two or three orders of magnitude greater than

their present concentration of 1.7 ppmv, in which

case, methane might well have been the dominant

greenhouse gas. With higher methane concentrations,

the number densities of hydrogen atoms in the upper

50 The Earth System

Geological evidence suggests that major

glaciations, extending all the way into the trop-

ics, occurred three times in Earth’s history:

the first around 2.2–2.4 billion years ago, con-

current with the rise of oxygen,

the second

between 600 and 750 million years ago, and the

latest so-called Permian Glaciation, 280 million

years ago. Until quite recently, most climate

researchers discounted this evidence on the

grounds that had the oceans ever been in a

completely ice-covered state, they would have

remained so. These skeptics argued that the

albedo of an ice-covered planet would be so low

that very little of the incident solar radiation

would be absorbed. Hence, the surface of the

Earth would have been so cold that the ice could

not have melted. Recently, this argument has

been called into question by proponents of the

following “freeze-fry scenario.”

i. During an unusually cold period, a

sufficiently large fraction of the Earth’s

surface becomes ice covered so that the

ice-albedo feedback mechanism described

in Section 10.3 renders the climate unstable:

the expanding ice cover cools the Earth’s

surface, the cooling causes the ice to expand

2.2 “Snowball Earth”

Continued on next page

It has been hypothesized that the decline in atmospheric methane concentrations brought about by the rise of oxygen might have

precipitated a sudden global cooling at this time.

Fig. 2.29 Artist’s conception of the onset of a world-

wide glaciation. [Courtesy of Richard Peltier.]

still farther, and the process continues until

the entire Earth becomes ice covered, as

depicted schematically in Fig. 2.29.

ii. During the ensuing snowball Earth phase,

the carbonate formation reaction (2.11)

cannot occur because the oceans are frozen.

P732951-Ch02.qxd 9/12/05 7:40 PM Page 50

2.5 A Brief History of Climate and the Earth System 51

atmosphere could have been orders of magnitude

higher than their current values. Under those condi-

tions, large numbers of hydrogen atoms would have

escaped to space, gradually raising the level of oxida-

tion of the Earth system. The abrupt transition of the

atmosphere from an anoxic state, with relatively high

concentrations of CH

,CO,H

S, and other reduced

gases, to a more oxidized state was marked by the

extinction of anaerobic life forms that are intolerant

of O

.The rise of O

set the stage for the evolution of

more complex life forms.

2.5.2 The Past 100 Million Years

During the Cretaceous epoch, which ended 65 mil-

lion years ago, surface air temperatures were sub-

stantially higher than they are today, especially at the

higher latitudes. This view is supported by the discov-

ery of remains of dinosaurs and lush tropical plants

dating back to that time in Siberia, Canada, and

other subarctic sites. Geological evidence indicates

that atmospheric concentrations of CO

were about

an order of magnitude higher at that time than they

are today. The Cretacous epoch was followed by an

extended interval of cooling and declining CO

con-

centrations, culminating in Pleistocene glaciation,

which began around 2.5 million years ago.

The cooling that set the stage for the Pleistocene

glaciation is widely attributed to the role of plate

tectonics in regulating the reactions (2.14) and (2.15)

in the carbonate–silicate cycle. Geological evidence

indicates that the rate of movement of the plates has

slowed down over the past 100 million years. A

reduced rate of ingestion of limestone sediments into

the mantle implies reduced metamorphism, which

would favor reduced volcanic emissions of CO

Meanwhile, the rise of the Himalayas following the

collision of the Indian and Asian plates (Fig. 2.30) is

believed to have increased the rate of weathering of

CaSiO

rocks, making more Ca

2

ions available for

However, continental drift continues and,

with it, metamorphism and volcanic

emissions of CO

into the atmosphere. In

the absence of a carbon sink, atmospheric

concentrations increase.

iii. Eventually, the combination of the increasing

greenhouse effect and the blackening of

parts of the ice sheets by windblown dust

raise the surface temperature up to the

threshold value at which ice in the tropical

oceans begins to melt. Once this process

begins, the ice-albedo feedback exerts a

powerful warming influence, which abruptly

flips the Earth system into an ice-free state.

iv. With the thawing of the oceans, carbonate

production resumes, but, compared to the

reaction time of the cryosphere in (iii), this

is a slow process, limited by the rate at which

weathering supplies Ca

2

ions. Hence, for

several million years, the Earth system

resides in the hothouse phase, with global-

mean temperatures initially as high as 50 °C.

The occurrence of an extended hothouse phase is

supported by the existence of rock formations

(i.e., banded iron formations and cap carbonates)

suggestive of a period of very high temperatures

around the time of the Permian glaciation.

2.2 Continued

Pleistocene (from the Greek: pleistos: most, and ceno: new) in reference to geological sediments. The Pleistocene epoch and the

subsequent Holocene (all new) epoch comprise the Quaternary period in the Earth’s history.

Fig. 2.30 Continental configuration 65 million years ago, at

the end of the Cretaceous epoch. Note the separation

between the Indian and Eurasian plates. [Courtesy of the U.S.

Geological Survey.]

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limestone formation in the oceans, thereby accelerat-

ing the removal of CO

from the atmosphere and

oceans. A decreasing source and increasing sink of

atmospheric CO

would account for the apparent

decline in atmospheric CO

levels, and the conse-

quent weakening of the greenhouse effect is consis-

tent with the observed cooling.

Another factor that contributed to the cooling was

glaciation of the Antarctic continent 15–30 million

years ago as it drifted into higher latitudes, which

would have increased the fraction of the incident

solar radiation reflected back to space. Other impor-

tant milestones in the drift of the continents toward

their present configuration were the opening of the

Drake passage in the southern hemisphere 15–30

million years ago, which gave rise to the formation of

the Antarctic circumpolar current, and the joining

of the North and South American continents at the

isthmus of Panama 3 million years ago. It has been

suggested that these events could have caused major

reorganizations of the oceanic thermohaline circula-

tion, resulting in a reduction of the poleward heat

flux in the North Atlantic, thereby accelerating the

cooling of the Arctic.

2.5.3 The Past Million Years

The past 2.5 million years have been marked by

climatic swings back and forth between extended

glacial epochs in which thick ice sheets covered

large areas of North America, northern Europe, and

Siberia and shorter interglacial epochs such as the

present one, in which only Antarctica (and some-

times Greenland) remain ice covered. Carbon and

oxygen isotopes (see Box 2.1) and the remains of

living organisms buried in marine sediments in

anoxic ocean basins around the world reveal a great

deal about the history of the pronounced climatic

swings during this so-called Quaternary period of

Earth’s history.

A more accurate and detailed history of the

climate swings of the past few hundred thousand

years, as revealed by an ice core extracted from

the dome of the Antarctic ice sheet, is shown in

Fig. 2.31. The temperature variations shown in

Fig. 2.31 are inferred from changes in concen-

trations of deuterium in the ice, as explained in

Box 2.1. Atmospheric concentrations of CO

and

are inferred from microscopic bubbles of air

that became trapped in the ice as the snow from

which it was formed was compressed and consoli-

dated. Because these gases tend to be well mixed,

their concentrations in the cores are indicative of

global conditions. The chemical signature of the

dust in the cores provides a basis for identifying it

with one or more specific source regions such as

the Gobi Desert, where fine particles of soil are

exposed directly to strong winds.

52 The Earth System

200

400

600

800

(pptv)

150

200

250

300

(ppmv)

0 50 100 150 200 250 300 350 400

−10

−5

Age (thousands of years)

Temperature change (°C)

Fig. 2.31 Comparison of methane, carbon dioxide, and estimated temperature (from oxygen and deuterium isotope ratios) from

the Vostok ice core, Antarctica, over the last 440 thousand years. The location of Vostok is indicated by the red dot in Fig. 2.13.

Note that the time axis runs from right to left. [Adapted from J. R. Petit et al., “Climate and atmospheric history of the past

420,000 years from the Vostok ice core, Antarctica.” Nature, 399, p. 431, 1999. Courtesy of Eric Steig.]

P732951-Ch02.qxd 9/12/05 7:40 PM Page 52

2.5 A Brief History of Climate and the Earth System 53

Based on the analysis of ice core records like

the ones shown in Fig. 2.31 and comparisons with

numerous (but less highly resolved and accurately

dated) marine sediment cores, it is clear that on

timescales of tens of thousands of years or longer,

temperature and a number of other climate-

related parameters vary coherently with one

another and that these variations are global in

extent. Global temperatures cooled at an irregular

rate during the extended glacial epochs and rose

much more rapidly at the beginning of the inter-

glacials, which have been recurring at intervals of

roughly 100,000 years. The last glacial maximum

occurred around 20,000 years ago. Atmospheric

and CH

concentrations have risen and fallen

synchronously with temperature, and the colder

epochs have been dustier, perhaps because of

higher wind speeds or drier conditions over the

source regions.

Figure 2.32 contrasts the extent of the northern

hemisphere continental ice sheets at the time of the

last glacial maximum with their current extent.

Parts of Canada were covered by ice as thick as

3 km. As a consequence of the large amount of

water sequestered in the ice sheets, the global sea

level was 125 m lower than it is today. The con-

centration of atmospheric CO

was 180 ppm, as

compared with a mean value of 260 ppm during

the current interglacial prior to the industrial

revolution. Hence, the Earth’s albedo must have

been higher at the time of the last glacial maximum

than it is today and the greenhouse effect must

have been weaker, both of which would have

favored lower surface temperatures. Temperatures

in Greenland were 10 °C lower than they are

today and tropical temperatures are estimated to

have been 4 °C lower.

There is evidence of a small time lag between

fluctuations in atmospheric CO

concentrations

and fluctuations in the volume of ice stored in the

continental ice sheets, with ice volume leading

. Hence, it would appear that the cause of these

fluctuations is intimately related to the growth and

shrinkage of the continental ice sheets. The CO

fluctuations represent a positive feedback that

amplifies the temperature contrasts between glacial

and interglacial epochs, as discussed further in

Section 10.3.2.

The pronounced climatic swings during the

Quaternary period are believed to be driven by sub-

tle variations in the Earth’s orbit that affect the sum-

mer insolation (i.e., the average intensity of incident

solar radiation) at high latitudes of the northern

hemisphere. During intervals when the summer inso-

lation is relatively weak, snow deposited during win-

ter does not completely melt, leaving a residual,

which, over a time span of thousands of years, accu-

mulates to form thick ice sheets. The high reflectivity

of the growing ice sheets exacerbates the coolness of

the summers, amplifying the orbital forcing. Based on

the same reasoning, it is believed that the continental

ice sheets are most prone to melting during periods

Fig. 2.32 The extent of the northern hemisphere continental ice sheets at the time of the last glacial maximum 20,000 years

ago (left) as compared with their current extent (right). [Courtesy of Camille Li.]

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