fundamentals of geobiology (knoll/fundamentals of geobiology) || geobiology of the phanerozoic

22
403 Fundamentals of Geobiology, First Edition. Edited by Andrew H. Knoll, Donald E. Canfield and Kurt O. Konhauser. © 2012 Blackwell Publishing Ltd. Published 2012 by Blackwell Publishing Ltd. 21 GEOBIOLOGY OF THE PHANEROZOIC Steven M. Stanley Department of Geology and Geophysics, University of Hawaii, 1680 East-West Road, Honolulu, Hawaii 96822, USA 21.1 The beginning of the Phanerozoic Eon Other chapters of this book trace important themes of geobiology through time. This chapter explores such themes as well, but it provides what amounts to a ‘hori- zontal’ rather than ‘vertical’ treatment, guiding the reader on a trip through time that focuses on groups of significant geobiologic events that occurred during par- ticular intervals of Earth’s history. In reviewing the history of Phanerozoic geobiology, it is appropriate to begin with the phenomenon that gave the Phanerozoic its name: the polyphyletic evolu- tion of skeletons that ushered in the Cambrian Period. After a long interval of ‘aragonite seas’ in the Proterozoic, a shift to ‘calcite seas’ came early in Cambrian time, when the molar Mg/Ca ratio of seawa- ter dropped below 2. Possibly the elevation of [Ca 2+ ] that contributed to this shift promoted the calcification of marine animals by increasing the supersaturation of seawater with respect to CaCO 3 (Brennan et al., 2004). There is evidence that the transition to calcite seas also led to the origins of skeletons consisting of low-Mg calcite, whereas the earliest Cambrian taxa produced skeletons of aragonitic or high-Mg calcite (Porter, 2007; Zhuravlev and Wood, 2008). The expansion of animal activity in the oceans had important consequences for marine geochemistry and sedimentology. For example, the polyphyletic produc- tion of skeletons in Early Cambrian time inevitably changed the CaCO 3 budget of the ocean. One result would have been a reduction of non-skeletal precipita- tion of CaCO 3 . Because silica occurs at a low concentra- tion in seawater, the advent of siliceous biomineralization must also have strongly affected the silica budget in the ocean (Maliva et al., 1989). In the absence of silica- secreting organisms, silica was relatively abundant in the ocean during Precambrian time, and as a conse- quence, early diagenetic cherts formed abundantly in peritidal marine sediments, possibly through microbial activity. Although demosponges, which produce spic- ules of silica, invaded offshore habitats early in the Paleozoic, they failed to suppress the precipitation of cherts in peritidal environments. On the other hand, the initial evolutionary radiation of the Radiolaria resulted in enough silica sequestration that cherts no longer formed in peritidal environments after Ordovician time. Cambrian strata typically exhibit low levels of biotur- bation (Droser and Bottjer, 1988). In the absence of heavy browsing by animals, microbial mats carpeted many areas of shallow Proterozoic seafloors. Thus, stromato- lites were widespread, as were ‘elephant skin’ sedimen- tary surfaces that formed when microbial mats crinkled. The advent of effective grazing in Cambrian time reduced the production of these structures (Garrett, 1970; Hagadorn and Bottjer, 1997). That microbial mats still formed sporadically in the Cambrian is indicated by the common occurrence of thrombolites. These are stromatolite-like forms that lack layering because of dis- ruption by burrowers or borers that did not exist before the Cambrian, or by obstructing seaweeds, which may also have been new on the scene (Aitkin, 1967; Grotzinger et al., 2005). Also present in Cambrian rocks are occa- sional stromatolites and flat-pebble conglomerates. The latter contain platy carbonate clasts that resulted from storm wave fracturing of well-laminated, often

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Page 1: Fundamentals of Geobiology (Knoll/Fundamentals of Geobiology) || Geobiology of the Phanerozoic

403

Fundamentals of Geobiology, First Edition. Edited by Andrew H. Knoll, Donald E. Canfield and Kurt O. Konhauser.

© 2012 Blackwell Publishing Ltd. Published 2012 by Blackwell Publishing Ltd.

21

GEOBIOLOGY OF THE

PHANEROZOICSteven M. Stanley

Department of Geology and Geophysics, University of Hawaii,

1680 East-West Road, Honolulu, Hawaii 96822, USA

21.1 The beginning of the Phanerozoic Eon

Other chapters of this book trace important themes of

geobiology through time. This chapter explores such

themes as well, but it provides what amounts to a ‘hori-

zontal’ rather than ‘vertical’ treatment, guiding the

reader on a trip through time that focuses on groups of

significant geobiologic events that occurred during par-

ticular intervals of Earth’s history.

In reviewing the history of Phanerozoic geobiology,

it is appropriate to begin with the phenomenon that

gave the Phanerozoic its name: the polyphyletic evolu-

tion of skeletons that ushered in the Cambrian Period.

After a long interval of ‘aragonite seas’ in the

Proterozoic, a shift to ‘calcite seas’ came early in

Cambrian time, when the molar Mg/Ca ratio of seawa-

ter dropped below 2. Possibly the elevation of [Ca2+]

that contributed to this shift promoted the calcification

of marine animals by increasing the supersaturation of

seawater with respect to CaCO3 (Brennan et al., 2004).

There is evidence that the transition to calcite seas also

led to the origins of skeletons consisting of low-Mg

calcite, whereas the earliest Cambrian taxa produced

skeletons of aragonitic or high-Mg calcite (Porter, 2007;

Zhuravlev and Wood, 2008).

The expansion of animal activity in the oceans had

important consequences for marine geochemistry and

sedimentology. For example, the polyphyletic produc-

tion of skeletons in Early Cambrian time inevitably

changed the CaCO3 budget of the ocean. One result

would have been a reduction of non-skeletal precipita-

tion of CaCO3. Because silica occurs at a low concentra-

tion in seawater, the advent of siliceous biomineralization

must also have strongly affected the silica budget in the

ocean (Maliva et  al., 1989). In the absence of silica-

secreting organisms, silica was relatively abundant in

the ocean during Precambrian time, and as a conse-

quence, early diagenetic cherts formed abundantly in

peritidal marine sediments, possibly through microbial

activity. Although demosponges, which produce spic-

ules of silica, invaded offshore habitats early in the

Paleozoic, they failed to suppress the precipitation of

cherts in peritidal environments. On the other hand, the

initial evolutionary radiation of the Radiolaria resulted

in enough silica sequestration that cherts no longer

formed in peritidal environments after Ordovician time.

Cambrian strata typically exhibit low levels of biotur-

bation (Droser and Bottjer, 1988). In the absence of heavy

browsing by animals, microbial mats carpeted many

areas of shallow Proterozoic seafloors. Thus, stromato-

lites were widespread, as were ‘elephant skin’ sedimen-

tary surfaces that formed when microbial mats crinkled.

The advent of effective grazing in Cambrian time

reduced the production of these structures (Garrett,

1970; Hagadorn and Bottjer, 1997). That microbial mats

still formed sporadically in the Cambrian is indicated by

the common occurrence of thrombolites. These are

stromatolite-like forms that lack layering because of dis-

ruption by burrowers or borers that did not exist before

the Cambrian, or by obstructing seaweeds, which may

also have been new on the scene (Aitkin, 1967; Grotzinger

et  al., 2005). Also present in Cambrian rocks are occa-

sional stromatolites and flat-pebble conglomerates. The

latter contain platy carbonate clasts that resulted from

storm wave fracturing of well-laminated, often

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404 Fundamentals of Geobiology

Age (Myr)

Neogene

Penn.

Miss.

Silu

r.P

aleo

c.C

reta

ceou

sJu

rass

icT

riass

icP

erm

ian

Dev

onia

nO

rdov

icia

nC

ambr

ian

Paleoc.

Eoc.

Early

Oligoc.

Late

Early

Late

Early

Middle

Late

Early

Late

Middle

Late

Early

MiddleLate

Early

Middle

Middle

Late

EarlyMiddle

Early

Middle

Late

50

100

150

200

250

300

350

400

450

500

F

J

K

M

R

O

B

A

G

C

H

I

L

P

D

N

Q

(o)?

Ordovician

Silurian

0 1 2 3 64 5 0–2 –1–3–4–5

Sedgwickiizone

Convolutuszone

–27–28–29–30–31

(n)

(m)

Llandov.

Wenl.

0 1 2 3 64 5–1 –2 –1–3–4–5–6

(l)Ludlow

Wenlock

–2–3–4–5–60 1 2 3 64 5 7

(d)

E. Trias.

M. Trias.

0 1 2 3 4–1 7 8

zoneYabeina

Neoschw.zone

EarlyGuad.

5 6 73 4

(f)

(h)

Serp.

Tournais.

Visean

21 222019 230 1 2 3 64 5

(p)

Mohawk

Chazyan

0 1 2 3 4–2 –1–3

Marj.

Delam.

E. Cam.

0 1 2–2 –1–3 –4–5–6–7–8–9–10–11

(r)Marj.

Sunwapt.

Stept.

0 1 2 3 4–2 –1–3 5

(q)δ13C

δ13C

δ13C

δ13C

δ13C

δ13C

δ13C

δ13C δ13C

δ13C

δ13C

δ13C

δ13C

δ13C

δ13C

δ13C

δ13C

δ13C

0 1 2 3 4–2 –1–3 –4–5–6–7–8–9–10

Silur.

Devon.(k)

(j)Givet.

Frasn.

Famenn.

–3–4–5–60 1 2 3

(i)

Mississippian

1 2 3 64 5 7 –2 –1–3–4–5–6–7–8 0

Famennian

E

Guad.

Loping.

0 1 2 3 4 5–2 –1–3 –3–4–5–6–7

(g)

(a)

34Ma

Eoc.

Olig.

0 1 2 0 1 2

33Ma

(b)

Paleoc.

Eoc.

0 1 2–1

57.2Ma

57.4Ma

0–2 –1

Perm.

Trias.

–2 –1–3–4–5–6–70 1 2 3 4–2 –1

(e)

–6–7–8–9–10

Stept.

Marj.

δ18O

δ18O

δ18O

δ18O

δ18O

δ18O

δ18O

δ18O

δ18O

δ18O

δ18O

δ18O δ18O

δ18O

(c)

Trias.

Jur.

0 1 2–2 –11 2 3 4 5

Figure 21.1 Stable isotope excursions that have been

documented in shallow marine strata in association with mass

extinctions. Eighteen intervals (A–R) contain a total of 26 such

δ13 C excursions. Corresponding to these, and trending in the

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Geobiology of the Phanerozoic 405

algal-bound, strata that in the absence of extensive

burrowing had been partly lithified by submarine

cementation (Sepkoski, 1982).

21.2 Cambrian mass extinctions

Several major extinctions occurred during the Cambrian

Period. The first came at the end of the Early Cambrian,

when redlichiid trilobites died out (Zhu et al., 2004), as

did nearly all archaeocyathid (sponge) reef builders

(Hill, 1972). The Delamaran/Marjuman stage boundary

of the Middle Cambrian also marks a major extinction of

trilobites, as do the Marjuman/Steptoean and Steptoan/

Sunwaptan stage boundaries of the Late Cambrian

(Palmer, 1998). It has been suggested that the three

Middle and Late Cambrian mass extinctions resulted

from episodic upward expansion of cold, poorly oxy-

genated waters that not only caused extinction of taxa in

shallow waters but also permitted the migration into

these waters of trilobites that had previously occupied

deeper habitats (Stitt, 1975; Palmer, 1984; Perfetta et al., 1999). While cooling may have caused these extinctions

of shallow-water taxa, it is unlikely that reduced oxygen

contributed because waters above wave base are always

oxygenated by the atmosphere.

The first three Cambrian mass extinctions illustrate a

pattern that characterizes major extinctions for the entire

Phanerozoic: they coincide with sharp excursions for

carbon and oxygen isotopes (shifts of δ13C and δ18O) for

skeletal carbonates (Fig. 21.1q, r). Numerous ad hoc

explanations have been offered to explain these various

excursions, nearly all quite reasonably focusing on one

or more factors that have changed the rate of burial of

organic carbon, which is isotopically light. It appears,

however, that a unifying explanation can largely account

for all of the excursions except a small number associ-

ated with global oceanic anoxia (Stanley, 2010). The

most important factor is the rate of respiration of bacte-

ria, which increases exponentially with temperature.

Because about 90% of carbon burial in the oceans takes

place along continental margins (Reimers et  al., 1992),

these locations are where climatic changes have their

greatest impact on bacterial respiration. When global

temperatures rise, so do bacterial respiration rates, and

therefore a larger proportion of carbon in particulate

organic matter is returned to the ocean in the form of

CO2 instead of being buried. The increased rate of

remineralization of isotopically light carbon results in a

global decline in δ13C for seawater. On the other hand,

when global temperatures fall, the rate of burial of

organic carbon rises and so does δ13C for seawater.

Significantly, δ18O in calcium carbonate follows the same

pattern, because of fractionation by organisms and also,

if glaciers expand, because the H2O containing the light

oxygen isotope, 16O, evaporates preferentially and is

preferentially locked up in glaciers. Every pair of global

carbon and oxygen isotope excursions coinciding with

a mass extinction has been either positive or negative,

reflecting global climate change (Fig. 21.1). Global cli-

mate change must have played a role in nearly all of

these mass extinction, the most significant of which will

be discussed below.

Three secondary climate-related aspects of the marine

ecosystem must also have contributed to the carbon iso-

tope excursions during times of global climate change

(Stanley, 2010): (1) growth or melting of clathrates

(icy materials along continental margins that contain

methane, which is isotopically very light carbon); (2) the

positive correlation between temperature and degree of

fractionation of carbon isotopes by phytoplankton,

although this relationship is weak at temperatures above

~15° C (Freeman and Hayes, 1992); (3) increased phyto-

plankton productivity during ‘icehouse’ conditions,

when strong latitudinal temperature gradients have

strengthened the upwelling of nutrient-rich waters.

The only conspicuous exceptions to the rule described

above for mass extinctions and stable isotopes are posi-

tive excursions for δ13C for intervals of global warming

such as the those of the Toarcian (Jurassic) and latest

Aptian and Cenomanian (Cretaceous), when a global

oceanic anoxia developed and huge amounts of isotopi-

cally light organic carbon were buried.

21.3 The terminal Ordovician mass extinction

Marine life diversified dramatically as the Ordovician

progressed, but then at the end of this period suffered one

of the largest mass extinctions of the Phanerozoic. This cri-

sis has been convincingly connected to a brief expansion

of continental glaciers in Gondwanaland, reflected by

tillites in many regions of Gondwanaland and what is

same direction, are 19 published δ18 O excursions, which are

displayed in the plots to the right of those depicting δ13 C.

Encircled letters on the left indicate temporal positions of

excursions. Blue indicates association with global cooling and

red, with global warming; black indicates absence of published

evidence of associated climate change. Horizontal scales

represent magnitudes of δ13 C and δ18 O excursions in ‰. Light

δ13 C in N is for organic carbon rather than carbonates, and

heavy δ18 O in H is for conodonts rather than bulk or skeletal

carbonate. Ordinates represent stratigraphic positions of

samples and are neither precisely linear with respect to time

nor scaled the same for all graphs (after Stanley, 2010).

Figure 21.1 Continued.

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406 Fundamentals of Geobiology

now southern Europe (review by Diaz-Martinez and

Grahn, 2007) and also by eustatic sealevel lowering,

documented in many stratigraphic sections around the

world. A strong positive shift for δ18O in the ocean

presumably resulted from both cooling and expansion of

glaciers and was paralleled by a shift for δ13C (Brenchley

et  al., 1994; Saltzman and Young, 2005) (Fig. 21.1o). The

extinction took place in two pulses: warm-adapted taxa

died out preferentially in the first pulse, which was the

larger of the two, as cold-adapted taxa migrated from

deep water and high latitudes to shallow seas positioned

at lower latitudes (Berry et al., 1995; Sheehan, 2001). These

patterns point to cooling, presumably associated with

increased seasonality, as the primary agent of extinction.

Cold-adapted taxa died out preferentially in the second

pulse, which took place at the very end of Ordovician

time as the ice age waned, perhaps 2 my after the first

pulse.

21.4 The impact of early land plants

The spread of early land plants during Silurian time

altered terrestrial landscapes, but vascular plants did

not appear until Late Silurian time, and not until late in

the Devonian did land plants first form forests. The evo-

lution of seeds near the end of Devonian time liberated

land plants from moist environments and thus added

another major step in the transformation of terrestrial

landscapes. This ecological expansion had two major

consequences for the physical environment. First, plants’

root systems stabilized river banks. Whereas braided

streams, which produced gravelly, cross-bedded depos-

its, prevailed on continents before the Devonian, mean-

dering rivers with firm banks first became widespread

during the Devonian, producing characteristic point bar

cycles with coarse (channel) sediment at the base and

fine (floodplain) sediment at the top. Second, the initial

global expansion of forests accelerated weathering

because the roots of land plants secrete acids and other

compounds that break down silicate minerals. Such

chemical weathering consumes CO2, and it appears that

accelerated weathering led to climatic cooling and conti-

nental glaciation in the Late Devonian through reduc-

tion of greenhouse warming (Retallack, 1997).

21.5 Silurian biotic crises

Each of four positive excursions for δ13C in Silurian

marine carbonates, the last at the very end of the period,

occurred immediately after a marine biotic crisis

(Saltzman, 2001, 2002). These excursions coincided with

glacial episodes and positive oxygen isotope excursions

(Loydell, 2007) (Fig. 21.1k–n). Thus, the Silurian crises

appear to have resulted at least in part from climatic

cooling.

21.6 Devonian mass extinctions

Three large mass extinctions struck during the Devonian.

The Givetian crisis, which marked the end of the Middle

Devonian, has been little studied, but it eliminated many

marine taxa, including numerous rugose coral families

(House, 2002). The Frasnian crisis of the Late Devonian

spanned perhaps 3 million years, nearly eliminating

the  previously flourishing coral–stromatoporoid reef

community (Copper, 2002). The Famennian crisis, which

was briefer but more severe, occurred at the end of the

Devonian, eliminating not only many invertebrate

marine taxa but also the heavily armored marine placo-

derm fishes and a variety of terrestrial plants. All three

Devonian biotic crises struck tropical taxa preferentially

and were associated with abrupt sea level declines and

positive δ13C and δ18O excursions in the ocean (Joachimski

and Buggisch, 2002; Buggisch and Joachimski, 2006)

(Fig. 21.1 j, k). Glacial deposits in eastern North America

confirm the expansion of glaciers in late Famennian

time (Brezinski et  al., 2008). Unlike the great terminal

Ordovician crisis, the Devonian mass extinctions mark-

edly restructured the marine ecosystem, in part by

destroying the coral-stromatoporoid reef community

(Droser et al., 1997).

21.7 Major changes of the global ecosystem in Carboniferous time

There was renewed continental glaciation during the

first (Tournasian) age of the Mississippian (early

Carboniferous) (Isaacson et  al., 2008), accompanied by

positive shifts for δ13C and δ18O (Fig. 21.1, h). Then

climates warmed. Coal swamps, colonized primarily by

lycopod plants and seed ferns, spread broadly over low-

land areas of the world early in Pennsylvanian (late

Carboniferous) time. Because anaerobic conditions and

tannic acid in these swamps excluded decomposing bac-

teria, reduced organic carbon was buried with little

decay. Furthermore, termites had not yet evolved, so

that, although subject to attack by fungi, trunks of dead

trees often fell into swamp waters largely intact

(Labandeira et al., 1997). As a result, a large amount of

organic carbon was buried, rather than being returned

to the atmosphere as CO2 via respiration by decompos-

ers. The consequent reduction of greenhouse warming

led to the largest glacial episode of the entire Phanerozoic,

with massive ice sheet growth in the Southern

Hemisphere. Thus, both δ13C and δ18O in the ocean

increased at the start the Serpukhovian, the last age of

the Mississippian (Fig. 21.1h). Also resulting from the

glacial expansion were eustatic sea level oscillations,

which produced the cyclical deposits on cratons

known  as cyclothems in North America and coal

measures in Europe.

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Geobiology of the Phanerozoic 407

Burial excludes organic matter from consumption by

aerobic consumers and bacteria, leaving behind in the

atmosphere oxygen that would otherwise have been

consumed in respiration. Therefore, the increased burial

of reduced carbon during the Carboniferous resulted in

a buildup of atmospheric oxygen. Today, oxygen consti-

tutes 21% of atmospheric oxygen, and it has been esti-

mated that this percentage rose to 35% during

Pennsylvanian time (Berner, 2006) (Fig. 21.2). This rise

appears to explain the evolution of giant insects, includ-

ing dragonflies with wingspans of 60 cm, during the

Pennsylvanian (insects’ assimilation of oxygen is lim-

ited by the absorption area of their spiracles) (Graham

et  al., 1995). High atmospheric oxygen levels probably

also increased the incidence of wildfires.

The late Paleozoic ice age not only had a biotic trig-

ger, but also major biotic consequences. Its initiation

resulted in a mass extinction near the end of

Mississippian time (the seventh largest such crisis of

the Phanerozoic Eon), and a new state of the marine

ecosystem. For every major marine taxon, rates of origi-

nation and extinction dropped at the start of this ice age

and remained low until its end (Fig. 21.3). This pattern

reflected the preferential loss of narrowly adapted trop-

ical taxa and the survival of taxa with broad thermal

tolerances that were resistant to extinction and that,

because of their widespread geographic distributions,

did not readily produce isolated populations that might

emerge as new species (Stanley and Powell, 2003;

Powell, 2005).

More generally, the reduction of atmospheric CO2 that

began in Devonian time and continued into the

Carboniferous altered the physiology of land plants.

Beerling and Berner (2005) concluded that a series of

feedbacks occurred. Stomata, the pores through which

gases pass to and from leaves, increase in density with a

decrease in atmospheric CO2 because more stomata are

needed for CO2 uptake. An increase in stomatal density

results in an increase in water loss. As atmospheric CO2

declined beginning in the Devonian, stomatal density

increased, and this would have increased heat loss from

leaves via evapotranspiration of water to the atmos-

phere. Large leaves, because of their low surface-to-

volume ratio, are prone to lethal overheating, and the

increased heat loss from leaves as CO2 declined appar-

ently permitted the increase in maximum leaf size that

has been documented for large plants during the

Devonian–Mississippian interval.

21.8 Low-elevation glaciation near the equator

A variety of evidence in the American Southwest indi-

cates that glaciers were well-developed at low eleva-

tions within about 8° of the equator in Late Pennsylvanian

and Early Permian time: a glaciated valley in the ances-

tral Rocky Mountains, diamictite containing striated

clasts, and widespread loessites (Soreghan et al., 2008).

The remarkable cooling of climates near the equator at

this time has yet to be explained, but it is certainly

35

30

25

20

15

10

5

0600 0100200300400500

Time (Millions of years ago)

Per

cent

age

of O

2 in

atm

osph

ere

Cam

bria

n

Ord

ovic

ian

Silu

rian

Dev

onia

n

Pen

nsyl

vani

an

Mis

siss

ippi

an

Per

mia

n

Tria

ssic

Jura

ssic

Cre

tace

ous

Cen

ozoi

c

Figure 21.2 Estimated changes in the volume of the oxygen

reservoir in the atmosphere during the Phanerozoic. (a)

Changes in the relative percentage of 13C in seawater, estimated

from the isotopic composition of limestones. (b) Estimated

changes in the portion of Earth’s atmosphere consisting of

free oxygen. Percentages for particular intervals are based on

estimates of the concentration of unoxidized carbon and sulfur

in sediments, the burial of which causes oxygen to build up in

the atmosphere. The broad band depicts uncertainties in

calculations. (A after Berner, 1987; B after Berner, 2006.)

400 350 300 250

Origination of genera

Terminal permianmass extinction

Guadalupianmass extinction

Ice age

Extinction of genera

Per

cent

age

chan

ge

Devonian PermianMississipp. Pennsylv.

Carboniferous

70

60

50

40

30

20

10

0

Ice age

Figure 21.3 Reduction of rates of origination and extinction of

marine genera at the start of the late Paleozoic ice age to their

lowest levels in all of Phanerozoic time. These rates returned

to normal levels precisely when the ice age ended, partway

through the Permian Period (after Stanley and Powell, 2003).

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408 Fundamentals of Geobiology

reminiscent of the so-called snowball Earth intervals of

the Proterozoic.

21.9 Drying of climates

The Permian period was marked by a drying of climates

on a global scale. This was at least partly a result of the

assembly of all continental regions of the world into the

supercontinent Pangaea: broad landlocked areas became

orographic deserts. This climatic change caused coal

swamps to shrink and seed plants, such as conifers, to

expand their ecological role. This floral transition actu-

ally began at high latitudes late in the Carboniferous

and did not reach the tropics until the early Permian

(DiMichele et  al., 2001). With the burial rate for wood

reduced, weathering (oxidation) of buried carbon

exposed by erosion eventually elevated the concentra-

tion of atmospheric CO2 to the degree that the ice age

ended (although a few small continental glaciers appar-

ently survived beyond the Sakmarian, the second age of

the Permian). Possibly, then, the end of the ice age had a

plate tectonic trigger.

21.10 A double mass extinction in the Permian

For many years it appeared that the crisis at the end of

the Permian, the largest mass extinction of all time, was

a protracted event. A number of patterns indicate that

there was actually a separate mass extinction at the end

of the penultimate (Guadalupian) age of the Permian

(Jin et al., 1994; Stanley and Yang, 1994), about 9 million

years before the terminal Permian event. For example,

all fusulinids that were relatively large or possessed a

honeycomb-like wall structure disappear at the end of

the Guadalupian Stage. These forms are just as preserv-

able as other fusulinids, so that the observed disappear-

ances clearly represent actual extinction.

Life on the land experienced two Permian transforma-

tions that coincided with those in the marine realm.

Therapsids (informally termed mammal-like reptiles,

although they were not reptiles) experienced two pulses

of extinction, and terrestrial floras simultaneously

underwent major changes (Retallack et al., 2006). During

the terminal Permian event, the Glossopteris flora of the

Southern Hemisphere died out, and the coal that it had

produced in moist environments ceased to form.

Coniferous floras also declined dramatically. Dicroidium,

a plant genus adapted to warm climates, spread pole-

ward. Terrestrial sediments indicate that climates in

many areas became drier, probably in part because

warmer temperatures elevated evaporation rates.

Marine deposits in Japan indicate that the ocean also

became increasingly stratified during the Permian.

A  block of Central Pacific seafloor that contains the

Permo-Triassic boundary was obducted onto the island

of Japan during the Jurassic (Isozaki, 1997). The

Guadalupian beds of this block consist of cherts formed

from radiolarian tests and stained red by ferric oxide

(Fig. 21.4). The deep sea throughout most of Guadalupian

time was obviously well oxygenated, presumably by

cold waters descending at the poles. There appear to

have been two phases of mass extinction during the

Guadalupian, the first entailed cooling associated with a

positive shift for marine δ13C (Fig. 21.1, f) (Isozaki et al., 2007). At the time of the second Guadalupian extinction,

Reappearanceof reefs

in Europe

Tria

ssic

Mid

dle

Ear

lyM

iddl

eLa

te

Gua

dalu

pian

Wuc

h.C

han.

Grie

s.D

ien.

Sm

i.S

pa.

Ani

sian

Ladi

nian

Per

mia

n

Gray anoxicchert

Siliceousshale

Carbonaceousshale

Siliceousshale

Gray anoxicchert

Red hematiticchert

Red hematiticchert

Sys

tem

se

ries

stag

e

Terminalpermian

extinction

Terminalguadalupian

extinction

AD

eep-

sea

anox

ia (

~20

mill

ion

year

s)

Sev

ere

anox

ia (

>10

mill

ion

year

s)Figure 21.4 Obducted rocks in Japan that illustrate the

episode of deep-water anoxia that took place in Late Permian

time. When anoxia began at the end of Guadalupian (Middle

Permian) time, gray chert replaced hematitic (highly oxidized)

red chert. An interval of severe anoxia, represented by even

darker sediments, began at the time of the terminal Permian

extinction. Deposition of hematitic chert resumed in Middle

Triassic time, and at this time reefs began to grow again in

shallow water (after Isosaki, 1997).

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Geobiology of the Phanerozoic 409

the deep sea sediments turned from red to gray,

indicating weaker oxygenation. Then, at the time of the

terminal Permian crisis, the these sediments turned

black, indicating that the ocean became highly stratified,

and respiration by aerobic bacteria soon eliminated free

oxygen in the deep sea. This pattern supports the

terrestrial evidence that climates twice became warmer

on a global scale, with the strongest pulse of warming

being associated with the terminal Permian crisis. Polar

regions became too warm to ventilate the deep sea with

cold downwelling waters. Negative δ13 C and δ18 O excur-

sions for shallow marine carbonates (Fig. 21.1.e and

21.1g) reflect these two steps of global warming.

Currently in favour is the idea that the volcanism that

produced the Siberian Traps – the largest continental

volcanic outpouring of the Phanerozoic – led to the

terminal Permian crisis. Many of the rocks thus

produced date to 251 Ma, the precise time of the mass

extinction. The lavas erupted through vast coal depos-

its, and it is thought that large quantities of CO2

suddenly entered Earth’s atmosphere not only from

Earth’s deep interior but also from the heating and

burning of coal, which would also have released meth-

ane. It has been suggested that a major volcanic episode

in China was similarly the ultimate cause of the

Guadalupian crisis. In any event, release of greenhouse

gases from coal, perhaps augmented by a submarine

release of methane hydrates, may have contributed to

the pronounced global shift toward isotopically light

carbon that is recorded in both marine and terrestrial

sediments at the time of the terminal Permian mass

extinction (Berner, 2002).

Three hypothesized kill mechanisms remain viable, at

least for some of the terminal Permian losses. Perhaps

CO2 that built up in the stagnant deep sea during the

Permian suddenly erupted to the surface, killing marine

life even in shallow water (Knoll et  al., 2007a). Or

Possibly hydrogen sulfide built up in the stagnant deep

sea and suddenly erupted (Kump et al., 2005). The sim-

plest idea is that the observed climatic warming – and

on the land the attendant increase in aridity – caused the

great mass extinction. Perhaps reflecting this agent of

extinction was the almost total destruction of low-

latitude floras (Rees, 2002), which may have been

subjected to lethally high temperatures.

21.11 The absence of recovery in the early Triassic

From beginning to end, the Early Triassic, which encom-

passed about 6 my, was characterized by highly reduced

terrestrial and marine biotic diversity. Aulochthonous

deep-sea deposits in Japan show that the deep sea

became well oxygenated again precisely at the end of

the Early Triassic, indicating the final return of climates

to something resembling their previous state (Isozaki,

1997). Nonetheless, there is evidence that pulses of

extinction, rather than a continued inhospitable state,

held back biotic recovery. Following the two negative

δ13C and δ18O excursions associated with the two

Permian mass extinctions, three similar excursions

occurred early in the Triassic, the last at the end of Early

Triassic time (Payne et  al., 2004) (Fig.  21.1d). Most

marine taxa recover so slowly from crises that their fos-

sil records have as yet failed to reveal pulses of Early

Triassic extinction, but the rapidly evolving ammonoids

and conodonts clearly experienced severe mass extinc-

tions that were more-or-less coincident with the carbon

isotope spikes, followed by rapid recoveries (Fig. 21.5)

(Stanley, 2009b). Comprehensive oxygen isotope analy-

ses have not yet been conducted for the Early Triassic,

but it is likely that the three mass extinctions were asso-

ciated with pulses of global warming.

21.12 The terminal Triassic crisis

The Triassic Period ended with one of the largest mass

extinctions of the Phanerozoic. The disappearance of

nearly all therapsids, which had benefited from an evo-

lutionary head start on the dinosaurs, permitted the lat-

ter to rise to dominance on the land (Benton, 1983; Olsen

et al, 2002). The terrestrial impact of the terminal Triassic

event is indicated by a sudden 60% reduction of pollen

species accompanied by a ‘spore spike,’ which probably

represented the opportunistic spread of ferns across

terrestrial habitats (Fowell et al., 1994).

The timing of Late Triassic extinctions has been con-

troversial, partly because of problematical stratigraphic

correlations. Nonetheless, it is evident that marine

extinctions occurred over a substantial interval of time

and that many marine taxa actually died out during or

at the end of the penultimate (Norian) age of the Late

Triassic rather than during the final (Rhaetian) stage

(review by Tanner et al., 2004).

The Triassic–Jurassic biotic transition on the land is

recorded by sediments in Eastern North America that

accumulated in rift basins produced in the early stages

of the breakup of Pangaea that created the Atlantic

Ocean. The Triassic ended very close to 200 Ma. Massive

volcanism took place at this time within the Central

Atlantic Magmatic Province (CAMP), spanning an inter-

val of perhaps 3 my (Marzoli et al., 1999; Knight et al., 2004; Whiteside et  al., 2007) (Fig. 21.6). Furthermore,

massive volcanism is indicated by an increase in the

osmium-187/osmium-188 ratio in marine mudrocks,

accompanied by an increase in the total abundance of

osmium and rhenium (Cohen and Coe, 2002). This syn-

chronicity has led to the suggestion that volcanic CO2

emissions triggered the terminal Triassic mass extinction

via greenhouse warming. It remains uncertain whether

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410 Fundamentals of Geobiology

the earlier (Norian) marine extinctions might also have

been associated with very early CAMP eruptions or had

some independent cause. In any event, stomatal densi-

ties provide independent evidence of global warming

across the Triassic–Jurassic boundary. They decreased

for fossil leaves from Greenland and Sweden, indicating

a mean annual temperature increase of 3–4 °C (Fig. 21.7);

simultaneously, average leaf width declined, presuma-

bly representing an adaptive shift that increased heat

loss and thus reduced thermal death of leaves exposed

to higher environmental temperatures (McElwain et al., 1999). As would be expected, negative δ13C and δ18O

excursions in the marine record (Fig. 21.1c) reflect this

global warming event. Possibly a sharp drop in atmos-

pheric pO2 coincident with the rise of CO

2 operated in

concert with global warming to cause extinctions of ter-

restrial vertebrates (Huey and Ward, 2005).

21.13 The rise of atmospheric oxygen since early in Triassic time

Falkowski et  al. (2005) have documented a general

secular increase over the past 205 my for δ13C in both

marine organic carbon and marine carbonates. They

have attributed this trend largely to an increase in the

biomass of marine phytoplankton – and, hence, in car-

bon burial – resulting from the diversification of coc-

colithophores and diatoms. They have also suggested

that an attendant increase in atmospheric pO2 permit-

ted the late Mesozoic appearance of placental mam-

mals, which require a high level of ambient O2 to

oxygenate embryos.

21.14 The Toarcian anoxic event

A subzone of the Toarcian (the final Jurassic stage), is

characterized globally by deep-marine organic-rich

black shales; at their base is a negative shift of δ13C, so

large (–6‰) that it has been thought necessarily to reflect

the release of methane hydrate from continental mar-

gins (Hesselbo et al., 2000; Beerling et al., 2002). Although

not on the scale of a major biotic crisis, heavy extinction

occurred at this time for marine taxa living in basins and

on continental shelves at depths greater than perhaps

50 meters (Jenkyns, 1988). Apparently, the oxygen mini-

mum layer rose to this general level, with lethal effects

on animal life. The Toarcian anoxic event spanned

perhaps only 200 000 years.

9

Age ( Ma)

Number of genera

AmmonoidsConodonts

Number of species

Columbites Zone

EG

LGD

ien.

Sm

ithia

nS

path

ian

251

250

249

248

247

246

245

Major conodontextinction

7

2

3

2

21

7

4

27

4

?

?

δ13C–2 –1 0 1 2 3 7 8

0 10 20 300 10 20 30 40 50

?

Carbonisotope

excursions

Ear

ly

tria

ssic

Figure 21.5 Similar patterns of radiation and mass extinction for early Triassic ammonoids and conodonts, with mass extinctions

coinciding with negative carbon isotope excursions. Numbers of species and genera are from global compilations (after Stanley, 2009b).

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Geobiology of the Phanerozoic 411

21.15 Phytoplankton, planktonic foraminifera, and the carbon cycle

The high abundance of C29

sterenes in Paleozoic organic

matter suggests that green algae played a larger plank-

tonic role in Paleozoic than post-Paleozoic seas (Knoll

et  al., 2007b). These ‘green plastid’ forms declined in

importance during the Mesozoic, while ‘red plastid’

phytoplankton (dinoflagellates, coccolithophores, and

diatoms) rose to dominance (Falkowski et al., 2004)

The coccolithophores arose in late Triassic time, but

only a single species is known to have survived into

Jurassic time, and then their diversity rose dramatically

until set back by the terminal Cretaceous mass extinc-

tion (Bown, 2005). During the Jurassic and Cretaceous,

detached coccoliths were a major component of pelagic

sediments. Favoured by the low Mg/Ca ratio and high

[Ca2+] of seawater, coccolithophores flourished espe-

cially in Cretaceous seas, forming the chalk that gave the

Cretaceous its name (Stanley et al., 2005).

Planktonic foraminiferans arose in Jurassic time and

diversified greatly during the Cretaceous, and they too

began to contribute considerable amounts of pelagic

carbonate sediment. A consequence of the expansion of

calcifying plankton was a huge increase in the conveyor-

belting of CaCO3 to subduction zones, where its burial

ultimately led to volcanic release of CO2. This release

has significantly supplemented the release of CO2 by the

metamorphism of shallow-water carbonates.

21.16 Diatoms and the silica cycle

Diatoms have become the most successful ‘red plastid’

phytoplankton group, in part because of their highly

efficient system for CO2 uptake, low quotas for trace

metals, and ability to store nutrients in a central vacuole

(Knoll et  al., 2007b). The diversification and ecological

expansion of marine diatoms during the Cretaceous

resulted in a reduction of the concentration of silica in

the ocean (Maliva et al., 1989).

21.17 Cretaceous climates

There has been much controversy about Cretaceous cli-

mates. Cool winter temperatures for the North Slope of

Alaska during the Cretaceous are indicated by the pres-

ence of dinosaurs, which were endothermic, and the

absence of reptiles, which are ectothermic (Clemens and

Nelmes, 1993). Nonetheless, terrestrial floras from the

North Slope indicate maximum summer temperatures

of ~13 °C and winter temperatures no lower than 2–8 °C

(Parrish and Spicer, 1988). It is universally agreed that

climates were warmest in Cenomanian–Turonian (mid-

Cretaceous) time, and floras close to the Arctic Ocean

indicate that this polar body of water was at or above

the freezing temperature of freshwater not only during

the Turonian but also during Coniacean (late Cretaceous)

time (Herman and Spicer, 1996).

It now appears that the global latitudinal temperature

gradient during much of Cretaceous time was fairly pro-

nounced, and yet the mean global temperature was

quite high during mid-Cretaceous (Cenomanian–

Turonian) time. Oxygen isotopes of marine fish teeth

and pristine (diagenetically unaltered) planktonic

foraminiferans indicate, respectively, for the mid-

Cretaceous shallow seas temperatures of ~32 °C and

28 °C in the tropics (compared to 24–28° C today) and

~25 °C and 20 °C at a paleolatitude of 40° (Pucéat et al., 2007). Tetraether lipids of marine Crenarchaeota

(prokaryotic plankton), which change their chemical

composition with temperature and are resistant to

Africa

HighAtlas

SouthAmerica

NorthAmerica

500 km

Volcanics Sills Dike swarms

Figure 21.6 The widespread distribution of igneous rocks of

the Central Atlantic Magmatic Province, which formed at the

end of Triassic time. Continental basalts of this province were

even more extensive than shown here because many have

been eroded away (after Marzoli et al., 1999).

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412 Fundamentals of Geobiology

diagenesis, indicate temperatures of 32–36° C for the

tropical Atlantic during the Cenomanian–Turonian

compared to 27–32 °C for the preceding, Albian, age

(Schouten et al., 2003).

21.17.1 Mid-Cretaceous anoxia

Massive eruptions of submarine lavas in the Pacific

Ocean began slightly before 125 Ma (late Barremian

time) and continued until ~80 Ma (mid-Campanian

time). These eruptions not only reduced the Mg/Ca ratio

of seawater, thus favouring the calcification of coccolith-

ophores and other taxa with calcitic, as opposed to arag-

onitic, skeletons, but they also sent a substantial amount

of CO2 into the atmosphere, accentuating greenhouse

warming. The Cenomanian–Turonian episode of

extreme global warming (~100–89 Ma) was in the mid-

dle of this interval. High rates of seafloor production

elevated sea level, and sluggish ocean circulation (an

absence of descending cold, oxygenating polar waters)

led to expansion of the oxygen minimum zone. Black

muds were deposited extensively even in relatively deep

waters of epicontinental seas (Larson, 1991) (Fig. 21.8).

The relative abundance in black shales of

2- methylhopanoids, which are membrane lipids found

in cyanobacteria and some other bacteria, indicate that

during major oceanic anoxic events of the Aptian and

Cenomanian, prokaryotes dominated many oceanic phy-

toplankton assemblages (Kuypers et al., 2004). Low δ15N

values for the organic matter in these black shales appar-

ently reflects a dominance of cyanobacterial nitrogen

fixation (air is characterized by light N). In contrast,

because the N/P ratio in the ocean was low and upwelling

was weak, eukaryotic phytoplankton were unable to

flourish.

Rates of extinction were elevated somewhat during

the Cenomanian–Turonian transition, a time of upward

expansion of dysaerobic waters (Leckie et al., 2002), but

losses did not rise to the level of a major crisis.

21.17.2 The puzzle of reef-building corals

A substantial contribution of corals to shallow-water

reefs in the tropics during Jurassic and early Cretaceous

time is puzzling for two reasons. First, this was an inter-

val of calcite seas, yet today corals produce aragonite.

Second, the concentration of atmospheric CO2 for this

interval was much higher than it is today (review by

Royer, 2003), and experiments have shown elevated

CO2  to have a negative effect on the calcification of

many modern coral species (Marubini and Thake, 1999;

Renegar and Riegl, 2005). Nonetheless, calculations

show that the high concentration of calcium in late

Mesozoic seawater may have compensated for the ele-

vated CO2, making the saturation state of CaCO

3 nearly

the same as today (Stanley et  al., 2005). Also, experi-

ments have shown that three species of modern corals

produce calcium carbonate consisting of about 30% cal-

cite in Cretaceous seawater (Ries et al., 2006); production

of such skeletal material in the late Mesozoic would

have enhanced coral skeletal growth. It is also possible

that late Mesozoic corals differed physiologically from

modern corals.

21.17.3 The terminal Cretaceous extraterrestrial event

The biotic crisis that brought the Mesozoic Era to an

end was only the fifth most destructive of the

Phanerozoic for marine life, but it has always been

granted special attention because of having eliminated

(a)

StomateStomate

Sweden8

Greenland

Triassic

(b)

Triassic JurassicJurassic

Tran

sitio

n

2

4

6

Mean annual temperature (°C above present level)

CO2 (multiple of present level)

Figure 21.7 Evidence from stomates of increases in atmospheric CO2 and mean annual temperature on Earth at the end of the

Triassic. (a) Illustration of stomatal cells in a leaf. (b) Increases in the proportion of stomata in fossil ginkgo and cycad leaves,

indicating a rise in atmospheric CO2 levels (after McElwain et al., 1999).

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Geobiology of the Phanerozoic 413

the dinosaurs. It also resulted in the sudden extinction

of a large percentage of gymnosperm and angiosperm

land plants from regions as far apart as North America

(Johnson and Hickey, 1990) and Japan (Saito et al., 1986),

and it resulted in the immediate ecological expansion of

ferns, as indicated by an abrupt decline of pollen and

increase of spores in terrestrial sediments (Tschudy and

Tschudy, 1986).

In 1980 the geologist Walter Alvarez, along with his

father, Luis (a Nobel Laureate in Physics), and Helen

Michel, announced the discovery of an iridium anom-

aly, a high concentration of the heavy metal iridium, at

the level of the terminal Cretaceous crisis. They recog-

nized this as an extraterrestrial signal because iridium is

very rare in Earth materials and relatively more abun-

dant in meteorites (Alvarez et al., 1980). Also soon dis-

covered at the level of the extinction were shocked

mineral grains, which are products of extraterrestrial

impacts on Earth (Bohor et  al., 1984); microtectites,

which are glassy spheroidal grains produced by the

rapid cooling of liquid droplets of materials blasted into

the atmosphere by an impact (Montonari et  al., 1983),

and minute diamonds, which can be produced only at

extremely high pressures (Carlisle and Braman, 1991).

The ultimate confirmation of the extraterrestrial

event – presumably an asteroid impact – at the end of

the Cretaceous stands as a major triumph for geology.

This was the discovery that the Chicxulub crater, which

borders Mexico’s Yucatan Peninsula and is imaged from

geophysical gravity data, formed at exactly the time of

the terminal Cretaceous mass extinction: igneous rocks

in the crater produced by the heat of the impact date

precisely to the Cretaceous–Paleocene transition

(Swisher et al., 1992).

A global negative excursion of δ13C in sediments at the

Cretaceous–Paleocene boundary has been taken to indi-

cate a collapse of phytoplankton productivity, and hence

a sharp reduction of light carbon burial in the deep sea.

Furthermore, carbon isotopic ratios ceased to display

the normal gradient from relatively high values for

planktonic taxa to relatively low values for deep-sea

benthos (D’Hondt et al., 1998) – a gradient reflecting the

preferential removal of δ12 C from the photic zone by

phytoplankton and transmission of isotopically light

carbon to the deep sea: here, too, is evidence of decreased

productivity by phytoplankton. Isotopic data from

deep-sea foraminifera indicate that recovery of biomass

by phytoplankton required about 3 my.

The immediate agent or agents of death in the termi-

nal Cretaceous crisis remain under debate. The area

where the asteroid struck contains large volumes of sul-

fate evaporites and limestones, which should have

released large amounts of SO2 and CO

2 at the time of

impact. These compounds, along with production of

nitric acid by heating of N2 and O

2 in the atmosphere,

would have adversely affected life by producing

PLI

MIO

OLI

G

EO

C

PA

L

MA

A

CM

PS

AN

C, T

CE

N

ALB

AP

T

BA

RH

AU

VA

LB

ER

TIT

H

100 65.5145CenozoicCretaceous

Oceaniccrust

production

Black shales

Long cretaceousnormal

35

Magnetic reversalsNormal

Reversed

30

25

Oce

anic

cru

st p

rodu

ctio

n (M

illio

ns o

fcu

bic

kilo

met

ers

per

mill

ion

year

s)

20

15

Time (Million years ago)

Figure 21.8 Black muds that became black shales accumulated in moderately deep waters in many regions during mid-Cretaceous

time, when there was also a high rate of oceanic crust production and an absence of magnetic reversals (after Larson, 1991).

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414 Fundamentals of Geobiology

strongly acid rain (D’Hondt et al., 1994). In addition, an

‘impact winter’ may immediately have developed, as

particles blasted into the atmosphere screened out the

sun’s rays (Pope et al., 1994). On the other hand, as these

particles descended to Earth, friction in the atmosphere

would have generated enormous heat (Melosh et  al., 1990). Support for dramatic, sudden warming at the end

of the Cretaceous comes from stomatal densities for

leaves, which indicate global warming by ~7.5 °C within

just 10 000 years (Beerling et al., 2002b).

21.17.5 The ascendancy of mammals and angiosperms: beneficiaries of the terminal Cretaceous crisis

Clearly the dinosaurs’ extinction opened the way for the

diversification of mammals. Mammals remained rela-

tively small in body size even in late Cretaceous time,

about 150 my after their origin. Dinosaurs had the jump

on mammals, however, having originated earlier in

Triassic time. Although the traditional view has been

that dinosaurs suppressed Mesozoic mammals via com-

petition, it is much more likely that the suppression was

via predation. Anyone who has seen the movie ‘Jurassic Park,’ with its reconstruction of the relatively small but

vicious predatory dinosaur Velociraptor, will appreciate

the victimization that mammals faced throughout

Mesozoic time. Supporting the idea the predation held

mammals back is the evidence that most Mesozoic

mammals had refugial life habits, many being small

burrowers or climbers or being active nocturnally.

More recently has it become evident that, following

the terminal Cretaceous mass extinction, terrestrial veg-

etation underwent a change paralleling that of terres-

trial quadrupeds. Although angiosperms (flowering

plants, including grasses and hardwood trees) experi-

enced considerable taxonomic diversification following

their mid-Cretaceous origin, their earliest representa-

tives were largely restricted to unstable habitats along

rivers (Doyle and Hickey, 1976). A flora well preserved

over a large area in central Wyoming by a sudden erup-

tion of volcanic ash suggests that gymnosperms and

spore plants dominated many undisturbed habitats

even in latest Cretaceous time (Wing et al., 1993). It was

not until the Paleocene that angiosperms first came to

dominate most terrestrial landscapes. Thus, the angio-

sperms, like the mammals, were serendipitous benefi-

ciaries of the meteorite impact that brought the Mesozoic

Era to a close.

21.18 The sudden Paleocene–Eocene climatic shift

Isotopic evidence from foraminiferans points to a dra-

matic change in the thermal structure of the ocean at the

very end of the Paleocene Epoch (Kennett and Stott,

1991). Throughout most of Paleocene time, oxygen

isotope ratios in foraminiferan skeletons were heavier

for deep-sea species than for shallow-water species,

indicating colder temperatures in the deep sea. At the

very end of Paleocene time, a dramatic shift occurred,

indicating that even close to Antarctica, the deep sea

suddenly warmed to temperatures close to those of sur-

face waters; cool, dense waters were no longer descend-

ing to the deep sea, and deep-sea foraminiferans suffered

mass extinction. It appears that at this time Earth experi-

enced a sudden pulse of global warming that lasted less

than 3000 years. At the same time δ13 C in soil organic

matter and skeletons of foraminiferans at all depths

in  the ocean experienced a sudden negative shift

(Magioncalda et al., 2004; Wing et al., 2005) (Fig. 21.1b).

This shift was so abrupt that some workers have attrib-

uted it at least in part to the release of isotopically light

carbon from methane hydrates along continental shelves

(Dickens et al., 1995; Kennett et al., 2003). Methane is a

powerful greenhouse gas, and although in about a dec-

ade it almost entirely oxidizes to form CO2, a weaker

greenhouse gas, if released over thousands of years at

the end of the Paleocene it could have substantially

enhanced greenhouse warming caused by elevation of

atmospheric pCO2. Also contributing to the negative

δ13 C shift would have been a positive feedback: the

increased rate of bacterial respiration along continental

margins (Stanley, 2010).

The magnesium content of calcite in planktonic

foraminiferans, which increases with temperature,

together with oxygen isotopes of this calcite, indicates a

sudden temperature increase of 4–5 °C in the tropical

Pacific Ocean at the end of the Paleocene (Zachos et al., 2003). Acidification of the ocean also occurred, with the

calcite compensation depth (the depth at which solution

of calcite begins) shoaling by more than 2 km; isotopic

evidence indicates that the thermal structure of the

ocean then recovered gradually during less than 50 000

years (Zachos et al., 2005).

Latest Paleocene floras of unique taxonomic composi-

tion have been discovered in Wyoming, and analysis of

their leaf morphologies has suggested that mean annual

temperature in this region increased by about 5 °C in less

than 10 000 years (Wing et al., 2005). In this same region,

mammalian faunas underwent major changes that

entailed the initial arrival from the Old World of artio-

dactyls and perissodactyls, the two major groups of

hoofed herbivores in the modern world (Clyde and

Gingerich, 1998).

21.18.1 A warm climate in the Eocene, but why?

Relatively warm climates persisted into the Eocene, well

after the terminal Paleocene warming subsided. It has

long been recognized that palm trees grew in Wyoming

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Geobiology of the Phanerozoic 415

during the Eocene and that alligators were able to exist

within the Arctic Circle. What remains to be determined

is to what extent greenhouse warming by high levels of

atmospheric CO2 was responsible for the persistence of

a remarkably warm global climate after the initial pulse

of global warming.

Analyses of stomatal densities on terrestrial ginkgo

leaves suggest that atmospheric CO2 levels were only

slightly above present levels during the Eocene (Royer

et  al., 2001). On the other hand, analyses of alkenones

produced by planktonic coccolithophores suggest that

atmospheric CO2 levels were about four times their

modern level (Pagani et al., 2005). Alkenones are carbon

compounds produced by coccolithophores that are

refractory to diagenesis, and coccolithophores fraction-

ate carbon isotopes of CO2 used in photosynthesis in a

manner that varies with the ambient concentration of

CO2, which reflects the atmospheric concentration of

CO2. The results of the alkenone analysis are likely to be

valid because they accord with independent evidence of

very warm Eocene climates.

The remarkably warm temperatures at high latitudes

during the Eocene appear to require a special explana-

tion. Extremely low δ18O values from Metasequoia wood

preserved at Axel Heiberg Island, inside the Arctic

Circle, apparently reflect transport of moisture north-

ward from the Pacific Coast of Mexico, and progressive

fractionation via loss of 18O through precipitation; this

northward flow of moist air would have transported

much heat (Jahren and Sternberg, 2002).

21.18.2 The origin of the modern climatic regime

Climates cooled on a global scale at the end of the

Eocene, as reflected in positive shifts of δ13 C and δ18 O in

the ocean (Fig. 21.1a). In many regions climates also

became drier, because cooler oceans contributed less

water to the atmosphere through evaporation. Terrestrial

floras first indicated this climatic shift. There is a linear

relationship between mean annual temperature and the

percentage of species in angiosperm floras that have

smooth-margined (as opposed to jagged-margined or

lobed) leaves. Although the slope of the leaf-margin

curve may have varied somewhat through time, any

substantial change in the percentage of smooth-margined

leaves in fossil floras provides a clear indication of a

change in mean annual temperature. A major decline in

this percentage took place in North America from the

Gulf Coast to Alaska at the end of the Eocene (Wolfe,

1971). Seeds from the London Clay of England indicate

that slightly earlier in the Eocene a similar transition

occurred from a flora resembling that of modern

Malaysia to a temperate flora (Collinson et al., 1981).

The Eocene–Oligocene transition ushered in the mod-

ern world, in which as climates became drier in many

regions, grasslands expanded at the expense of forests.

(Trees require a consistent supply of water, whereas

grass taxa typically tolerate seasonal drought.)

A mammalian fauna of the Mongolian Plateau

records the replacement of forested habitats by open

habitats during the Eocene–Oligocene transition

(Meng and McKenna, 1998). Many medium-sized

hoofed animals, which are most common in forested

habitats, disappeared, as did tree-climbing taxa such

as primates. At the same time, species of rodents, rab-

bits, and open-country taxa with teeth adapted for

feeding on harsh grasses appeared, along with large

herbivores having the stamina to outrun predators in

open terrain.

In the marine realm, a second-order mass extinction

occurred in Late Eocene time, with a preferential loss of

warm-adapted molluscan taxa (Hansen, 1987; Hickman,

2003). Oxygen isotopes of mollusks and fish otoliths (ear

bones) from coastal plain deposits of the Mississippi

Embayment indicate a decline from tropical temperatures

between early Eocene and early Oligocene time, with

a  winter reduction of ~5 °C and a summer reduction

of  ~3 °C (Kobashi et  al., 2001, 2004). In other words,

the  climate change entailed an increase in seasonality.

Similarly, from the late Eocene into the Oligocene, radio-

larians in the equatorial Pacific experienced numerous

extinctions of purely tropical species and an increase in

cosmopolitan taxa with relatively broad thermal adapta-

tions (Funakawa et al., 2006). Planktonic foraminiferans

underwent stepwise extinction during the same interval

(Keller, 1983).

21.19 The cause of the Eocene–Oligocene climatic shift

Traditionally the Eocene–Oligocene climatic shift has

been attributed to the formation of the Circumantarctic

Current (Kennett et  al., 1975). This current traps water

that consequently becomes very cold. In the present

ocean, the relatively high density of this cold water,

enhanced by an elevation of salinity through sea ice for-

mation, causes downward convection, producing the

cold bottom layer of the ocean by spreading to the far

north in both the Atlantic and Pacific. The Circumantarctic

Current formed when Antarctica became isolated over

the South Pole as South America and Australia broke

away from it. Thus, the Drake Passage and Tasmanian

Gateway formed – and with them the modern polar

gyre came into being. Upward mixing of cold, deep

waters that formed in the vicinity of Antarctica would

have cooled climates throughout the globe.

Coincidentally, at ~35 Ma (close to the Eocene–

Oligocene transition) the tectonic deepening of the

Greenland-Iceland-Faeroes Ridge permitted downward

cold-water convection in the North Atlantic (Davies

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416 Fundamentals of Geobiology

et  al., 2001). Today, the water descending in the North

Atlantic spreads throughout the ocean above dense

Antarctic bottom water. A portion of the present

Antarctic ice sheet had formed and was producing gla-

cial marine deposits by the start of the Oligocene (Ivany

et al., 2006). In the north, ice-rafted debris began to reach

the Norwegian-Greenland Sea by 38 Ma, meaning that

at least some isolated glaciers had formed by this time

on Greenland (Eldrett et al., 2007).

A tectonic evaluation of the opening of the Drake

Passage, based on seismology, indicates that the pas-

sage began to form during middle Eocene time (Eagles

et al., 2006). In addition, a variety of evidence indicates

that the Tasmanian Gateway began to form slightly

before the end of Eocene time (~Ma) (Stickley et  al., 2004). The implication is that the Circumantarctic

Current arose during the latter part of the Eocene.

Neodymium isotopes in fossil fish teeth support this

timing (Scher and Martin, 2006). The 143Ne/144Ne ratio

has long been higher in the Pacific than in the Atlantic,

reflecting circumpacific volcanism, and yet with mixing

between the two oceans today there is only a small dif-

ference between them in this ratio. Early in the Cenozoic,

the difference was much larger, but during the middle

Eocene, apparently in response to the formation of an

incipient Circumantarctic Current, this ratio diminished

dramatically.

Alternatively, it has been suggested that a decrease in

greenhouse warming, via lowering of atmospheric

pCO2, produced the Eocene–Oligocene climatic change.

However, alkenones in coccolithophores appear to indi-

cate that although pCO2 dropped in mid-Cenozoic time,

it did not do so until ~32 Ma, some 2 my after the global

climatic change occurred (Pagani et al., 2005) (Fig. 21.9).

Possibly the timing of this pCO2 decline will be revised

in the future.

21.20 The re-expansion of reefs during Oligocene time

Despite heavy losses in the terminal Cretaceous mass

extinction, reef-building scleractinian corals retained

substantial taxonomic diversity at the start of the

Cenozoic. They nonetheless produced very few reefs of

any size during the Paleocene or Eocene. Something

prevented scleractinians from flourishing until

Oligocene time. Three possibilities are evident:

1 The chemistry of the oceans shifted from calcite to

aragonite seas close to the Eocene-Oligocene transition,

favouring calcification by scleractinians (Stanley and

Hardie, 1998).

2 Atmospheric CO2 declined markedly during early

Oligocene time (Pagani et al., 2005), and this favored the

precipitation of calcium carbonate in the ocean.

3 Possibly until Oligocene time Cenozoic corals lacked

the symbiotic algae that today promote their calcifica-

tion. The molecular clock indicates that the symbiotic

algae of modern reef-building corals originated during

the Eocene (Pochon et al., 2006). The implication is that

the terminal Cretaceous mass extinction eliminated

more ancient types of symbiotic algae in reef-building

corals, and corals were not recolonized by algae until at

least Eocene time.

21.21 Drier climates and cascading evolutionary radiations on the land

The fossil record of phytoliths, silica bodies secreted by

plants, indicates that the modern taxa of grasses adapted

to open habitats diversified in late Oligocene and early

Miocene time (Strömberg, 2004). The expansion of open

habitats that began at this time produced cascading evo-

lutionary radiations of plant and animal taxa adapted to

these habitats and led to the high diversity of these taxa

in the present world (Stanley, 1990) (Fig. 21.10). Not only

have grasses diversified since this time, but also weeds

(the family Compositae), which opportunistically

occupy open spaces in grasslands. In addition, the

Muridae (Old World rats and mice) and songbirds, both

pC

O2

(ppm

)

Time (Million years ago)

4050 30 20 10 0

OligoceneEocene Miocene

1000

1500

2000

2500

500

Figure 21.9 Estimates of the concentration of CO2 in Earth’s

atmosphere from Eocene through Miocene time, based on the

carbon isotopic composition of alkenones in calcareous

nanoplankton. The top of the stippled band represents the

maximum estimate, the bottom of this band represents an

intermediate estimate, and the dashed line represents a

minimum estimate. This analysis indicates that the level of

atmospheric CO2 was very high in the Eocene and earliest

Oligocene and began to drop precipitously about 32 million

years ago (vertical red line), some 2 million years after glaciers

expanded in Antarctica and climates changed throughout the

world (after Pagani et al., 2005).

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Geobiology of the Phanerozoic 417

of which contain many species that feed on seeds of

grasses and weeds, began spectacular evolutionary

radiations. Finally, this was the time when the radiation

of the snake family Colubridae began. Snakes can slither

along branches to consume songbirds’ eggs and chicks

and can make their way down small rodent holes. The

family Colubridae, which contains most species of mod-

ern snakes that are not constrictors and includes all ven-

omous forms, arose and began a spectacular evolutionary

radiation in the Miocene.

21.21.1 Climate change, extinction, and the spread of C4 grasses

There is evidence of widespread aridification related

to cooling at ~7–6 Ma (close to the end of the Miocene).

The global volume of glacial ice increased, causing the

Messinian sea level fall of at least 30 m (Aharon et  al., 1993). At the same time, the oceans cooled at high and

middle latitudes in both hemispheres (Poore and

Berggren, 1975) and grasslands replaced woodlands in

many regions (Webb, 1977; Bernor et  al., 1996; Gentry

and Heizmann, 1996). The largest extinction event of the

past 30 million years for North American mammals

occurred at this time, largely in response to the spread of

grasslands (Webb, 1984).

Carbon isotope ratios of the teeth of herbivores

reflect the isotope ratios of the food that they eat,

although fractionation occurs as the food is assimi-

lated. A marked increase in carbon δ13 C occurred on a

global scale in mammal teeth preserved in sediments

ranging from ~7 to 6 Ma (Cerling et al., 1993) (Fig. 21.11).

This change reflected the worldwide spread of C4

grasses, a group that utilizes a different photosynthetic

pathway than C3 grasses and fractionates carbon iso-

topes in such a way that their tissues contain a higher

percentage of 13C. Warm, seasonally dry savannah hab-

itats favour C4 grasses. In contrast, C3 grasses require

perennial moisture in the tropics or a cool, moist grow-

ing season in nontropical regions; Mediterranean

climates and northern temperate climates provide the

latter conditions.

Grasses(Gramineae)

Herbs and weeds(Compositae)

Old worldrats and mice

(Muridae)

Modernsongbirds

Modern snakes(Colubridae)

Time(Million years

ago)

5

10

15

20

25

Consumers

500 species

Producers

10,000 species

Figure 21.10 Cascading evolutionary radiations of

terrestrial taxa during the past 25 million years. Grasses

and weeds diversified dramatically, and their prolific

production of seeds was partly responsible for a great

expansion of rats, mice, and songbirds. Colubrid snakes,

which feed on rats and mice and the eggs and chicks of

songbirds also experienced a remarkable radiaton (after

Stanley, 1990).

20

15

10

5

0

d13C

–15 –10 –5 0 5

Tim

e (M

illio

n ye

ars

ago)

Figure 21.11 Major shifts in carbon isotopes between

7 million and 6 million years ago, indicating the spread of

C4 grasses. The plotted values are carbon isotope ratios from

ancient soils and mammal teeth from Pakistan and North

America (after Cerling et al., 1993).

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418 Fundamentals of Geobiology

Because grasses contain abundant phytoliths (silica

bodies) that wear down the teeth of grazers, grazing

mammals generally have molars that are initially taller

than those of mammals that browse on softer leafy veg-

etation. During the Miocene, hypsodont horse species

(ones with tall molars) and then ones classified as ‘very

hypsodont’ increased in numbers, while mesodont

species (those with medium-tall molars) declined; after

extinction of the last North American mesodont forms

at 11–12 Ma, only species adapted for grazing remained

(Hulbert, 1993). This net trend developed as a result of

the expansion of grasslands during the Miocene (Webb,

1984). Then there was an abrupt shift for American

horses toward very hypsodont teeth at 7–6 Ma at the

time when C4 grasses proliferated. Because C4 grasses

contain on average about five times as many phytoliths

per volume of tissue as C3 grasses, it seems evident

that very tall teeth were at a premium for horses that

fed on C4 grasses (Stanley, 2009a, p. 458). Horses

employ inefficient hind-gut digestion and are therefore

required to feed for more hours every day than other

large herbivores. Presumably, North American horse

species that lacked very hypsodont molars experienced

shortened lifespans as C4 grasses expanded, and their

overall birth rates declined to levels that could not

sustain populations.

21.21.2 The initiation of the modern ice age

The modern ice age of the Northern Hemisphere, dur-

ing which we still live, cannot be attributed to green-

house cooling, because all indications are that there was

no decline in atmospheric pCO2 during the onset of the

ice age, between 3.5 and 3.0 my ago. At this time plate

tectonic movements emplaced the Isthmus of Panama

between North and South America. The Arctic region

today is cold because the Arctic Ocean is isolated, with

little inflow of warm waters from the Pacific or Atlantic

oceans. The most important factor here is that north-

ward-flowing Atlantic waters today sink north of

Iceland, depriving the Arctic Ocean of their warmth.

These waters descend not only because they become

cold, but also because they are unusually saline as a

result of high evaporation rates farther south, in the

trade wind belt.

Before the emplacement of the Isthmus of Panama,

tropical Atlantic waters would have flowed into the

Pacific and compensatory flow would have moved

water in the opposite direction. Therefore, the Atlantic

would have been less saline than it is today. It follows

that Atlantic waters would have sunk farther north than

they do today, i.e. within the Arctic Ocean. With the

inflow of these warm waters, the Arctic would have

been much warmer then than it is today. If these infer-

ences are correct, the formation of the Isthmus of Panama

may have caused the modern ice age of the Northern

Hemisphere by isolating the Arctic Ocean from warm

Atlantic waters (Stanley, 1995).

21.21.3 Biotic consequences of the modern ice age of the Northern Hemisphere

Cooling of the ocean during the modern ice age of the

Northern Hemisphere caused extinctions of marine life

in the North Atlantic region, where temperature declines

were concentrated (Raffi et  al., 1985; Stanley, 1986).

Elsewhere, extinctions were minor, partly because of the

geographic pattern of cooling and partly because Earth

had already moved into a glacial mode at the end of the

Eocene, so that many biotas were already adapted to

cool mean annual temperatures and pronounced sea-

sonality (Stanley, 1986).

Because surface waters of the ocean cooled and

yielded less moisture to the atmosphere late in the

Pliocene, climates became drier in many regions. In

Africa, the result was a contraction of forests and a

diversification of antelopes adapted to open habitats

(Vrba, 1985). In portions of South America, rainforests

also contracted (Hooghiemstra and van der Hammen,

1998).

Human evolution was strongly impacted by vegeta-

tional changes in Africa (Stanley, 1992). Australopithecus,

which was ancestral to the human genus Homo, undoubt-

edly spent considerable time on the ground, but it also

possessed numerous adaptations for climbing trees:

upward-directed shoulder sockets and also long arms,

long fingers, and long toes with the ability to grasp. For

animals having a brain size little above that of a chim-

panzee and a capacity for only slow movement on the

ground, these adaptations were necessary for tree-

climbing to avoid vicious African predators. Woodlands

contracted in Africa at about 2.5 Ma, when the northern

ice age intensified, and Australopithecus, which was

dependent on woodlands, died out.

At about the time when Australopithecus died out,

Homo, the modern human genus, came into being. The

large brain of Homo results from delayed development.

Monkeys, apes, and humans in the womb grow brains

that are about 10% of an embryo’s body size. For mon-

keys and apes, brain size relative to body size falls back

after birth, but for humans the 10% relationship persists

for about a year after birth. We humans experience a

general delay in development that endows us with most

of our adult brain size (at an age of about 1 year, we

assume the apes’ slower postnatal rate of brain growth).

Delayed development also saddles humans with rela-

tively immature, helpless infants. Adult and neonate

body and brain sizes indicated by fossils show that early

Homo possessed our delayed development but

Australopithecus did not (Stanley, 1992). In contrast to

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Geobiology of the Phanerozoic 419

Australopithecus females, early females of our genus

could not have climbed trees habitually because they

could not have climbed with one arm while carrying a

helpless infant in the other. Therefore, Homo must have

evolved from a population of Australopithecus confined

to the ground, perhaps after nearly all other populations

died out as forests shrank in Africa. We can be thankful

that this population happened to survive by evolving

into our genus, whose intelligence was undoubtedly

essential to its survival amidst many large predatory

mammals.

21.21.4 The younger Dryas: evidence for an exraterrestrial impact

About 12 900 years ago, as the Northern Hemisphere

was emerging from its most recent glacial maximum,

glaciers suddenly re-expanded; they shrank back again,

permanently, about 1300 years later. The interval before

the glaciers re-expanded is known as the Younger Dryas.

Numerous species of large mammals died out in North

America at the start of the Younger Dryas, and the Clovis

human culture also disappeared. There is now evidence

that a comet may have struck North America, producing

the Younger Dryas and associated events (Firestone

et  al., 2007). At numerous sites across North America,

there is a black layer (usually less than 3 cm thick)

precisely at the level of the mammalian extinctions and

termination of the Clovis culture. Containing charcoal,

soot, glassy carbon, and carbon spherules, this layer

represents widespread burning (Fig. 21.11). At the

Murray Hills site in Arizona, the black layer fills mam-

moth footprints and is draped over a Clovis fireplace

and nearly complete mammoth skeleton. The black

layer also contains helium, trapped within cagelike

fullerene molecules, that has an extraterrestrial isotopic

signature (Fig. 21.12e). Also present are magnetite parti-

cles. A high concentration of iridium has been reported

at some sites, but some researchers (e.g. Paquay et  al., 2009) have failed to duplicate the results. These various

materials may have come from the core inside the icy

exterior of a comet. The compositions of only a few com-

ets have been analysed. They vary considerably in com-

position and differ from chondritic meteorites, for

example in lacking nickel (Lisse et al., 2007). The most

compelling evidence for an extraterrestrial impact is

supplied by nanodiamonds that are of a type considered

to result only from impact events and that occur in vast

numbers at many sites where additional evidence is

found (Kennett et al., 2009).

A comet has been favoured as the agent of this crisis

because shocked minerals and microtectites – signatures

of a meteorite impact – are absent. To start numerous

fires, a comet would have had to fragment before land-

ing. This would not have been unlikely, because comets

are weak bodies, containing about 75% pore space. A

large peak in the abundance of ammonia and nitrate in a

Greenland ice core dates to 12 900 years ago (dated

by  counting ice varves), and their protracted decline

indicates the occurrence of widespread wildfires for

50 years. How an impact event may have produced the

Younger Dryas remains to be more fully explored.

200 mm

(b)

200 mm

(c)

(e)

(a)

5 mm

(d)

150 mm

Figure 21.12 Materials found in the YDB layer at many

Clovis sites. (a) Soot particles. (b) and (c) Exterior and

cross-sectional views of low-density carbon grains. (d)

Magnetic microspherule. (e) A model of a fullerene,

which is a cage-like molecule formed of carbon atoms; a

helium atom is shown entering the fullerene molecule on the

left, and one is trapped in the fullerene molecule on the right

(after Firestone et al., 2007).

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420 Fundamentals of Geobiology

Through the extinction only 12 900 years ago of

numerous North American mammals – among them

three elephant species, a sabertooth cat, a fast-running

bear that stood six feet tall at the shoulder, and a beaver

the size of a black bear – we modern humans have inher-

ited a highly impoverished land mammal fauna. This

situation should add value to the species that remain.

References

Aharon P, Goldstein SL, Wheeler CW, Jacobson G (1993) Sea-

level events in the South Pacific linked with the Messinian

salinity crisis. Geology 21, 771–775.

Aitken JD (1967) Classification and environmental significance

of cryptalgal limestones and dolomites, with illustrations

from the Cambrian and Ordovician of southwestern Alberta.

Journal of Sedimentary Petrology 37, 1163–1178.

Alvarez LW, Alvarez W, Asaro F, Michel HV (1980)

Extraterrestrial cause for the Cretaceous-Tertiary extinction.

Science, USA 208, 1095–1108.

Beerling DJ, Berner RA (2005) Feedbacks and the coevolution

of plants and atmospheric CO2. Proceedings of the National

Academy of Sciences 102, 1302–1305.

Beerling DJ, Lomas MR, Groecke DR (2002a) On the nature of

methane gas-hydrate dissociation during the Toarcian and

Aptian ocean anoxic events. American Journal of Science 302,

28–49.

Beerling DJ, Lomax BH, Royer DL, Upchurch GR, Kump LR

(2002b) An atmospheric pCO2 reconstruction across the

Cretaceous-Tertiary boundary. Proceedings of the National Academy of Sciences, USA 99, 7836–7840.

Benton MJ (1983) Dinosaur success in the Triassic; a noncom-

petitive ecological model. The Quarterly Review of Biology 58,

29–55.

Berner RA (1987) Models for carbon and sulfur cycles and

atmospheric oxygen: Application to Paleozoic geologic his-

tory. American Journal of Science 287, 177–196.

Berner RA (2002) Examination of hypotheses for the Permo-

Triassic boundary extinction by carbon cycle modeling. Pro-ceedings of the National Academy of Sciences, USA 99, 4172–4177.

Berner RA (2006) GEOCARBSULF: A combined model for

Phanerozoic atmospheric 02 and C02. Geochimica et Cosmochimica Acta 70, 5653–5664.

Bernor RL, Fahlbusch V, Andrews P, et al. (1996) The evolu-

tion of western Eurasian Neogene mammal faunas: A

chronologic, systematic, biogeographic, and paleoenviron-

mental synthesis. In: The Evolution of Western Eurasian Neogene Mammal Faunas (eds Bernor RL, Fahlbusch V,

Mittman H-W). Columbia University Press, New York,

pp. 449–459.

Berry WBN, Quinby-Hunt MS, Wilde P (1995) Impact of Late

Ordovician glaciation-deglaciation on marine life. In: Effects of Past Global Change on Life. Studies in Geophysics. National

Academy Press, Washington, DC, pp. 34–46.

Bohor BF, Foord EE, Modreski PJ, Triplehorn DM (1984)

Mineralogic evidence for an impact event at the Cretaceous-

Tertiary boundary. Science 224, 867–869.

Bond DPR, Wignall PB, Wang W, et  al. (2010) The mid-

Capitanian (Middle Permian) mass extinction and carbon

isotope record of South China. Palaeogeography, Palaeoclimatology, Palaeoecology 292: 282–294.

Bown PR (2005) Selective calcareous nannoplankton survi-

vorship at the Cretaceous-Tertiary boundary. Geology 33,

653–656.

Brenchley PJ, Marshall JD, Carden GAF, et  al. (1994)

Bathymetric and isotopic evidence for a short-lived Late

Ordovician glaciation in a greenhouse period. Geology 22,

295–298.

Brennan S, Lowenstein TC, Horita J (2004) Seawater chemistry

and the advent of biocalcification. Geology, 32 473–476.

Brezinski DK, Cecil CB, Skema VW, Stamm R (2008) Late

Devonian glacial deposits from the eastern United States sig-

nal an end of the mid-Paleozoic warm period. Palaeogeography, Palaeoclimatology, Palaeoecology 268, 143–151.

Buggisch W, Joachimski MM (2006) Carbon isotope stratigra-

phy of the Devonian of central and southern Europe.

Palaeogeography, Palaeoclimatology, Palaeoecology 240, 68–88.

Carlisle DB, Braman DR (1991) Nanometre-size diamonds in

the Cretaceous/Tertiary boundary clay of Alberta. Nature

352, 708–709.

Cerling TE, Wang Y, Quade J (1993) Expansion of C4 ecosys-

tems as an indicator of global ecological change in the late

Miocene. Nature 361, 344–345.

Clemens WA, Nelms LG (1993) Paleoecological implications of

Alaskan terrestrial vertebrate fauna in latest Cretaceous time

at high paleolatitudes. Geology 21, 503–506.

Clyde WC, Gingerich PD (1998) Mammalian community

response to the latest Paleocene thermal maximum; an iso-

taphonomic study in the northern Bighorn Basin, Wyoming.

Geology 26, 1011–1014.

Cohen AS, Coe AL (2002) New geochemical evidence for the

onset of volcanism in the Central Atlantic magmatic prov-

ince and environmental change at the Triassic–Jurassic

boundary. Geology 30, 267–270.

Collinson ME, Fowler K, Boulter MC (1981) Floristic changes

indicate a cooling climate in the Eocene of southern England.

Nature 291, 315–317.

Copper P (2002) Silurian and Devonian reefs: 80 million years

of global greenhouse between two ice ages. Society for Sedimentary Geology Special Publication 72, 181–238.

Davies R, Cartwright J, Pike J, Line C (2001) Early Oligocene

initiation of North Atlantic Deep Water formation. Nature

410, 917–920.

D’Hondt S, Pilson MEQ, Sigurdsson H, Hanson AK, Jr., Carey S

(1994) Surface-water acidification and extinction at the

Cretaceous–Tertiary boundary. Geology 22, 983–986.

D’Hondt S, Donaghay P, Zachos JC, Luttenberg D, Lindinger

M (1998) Organic carbon fluxes and ecological recovery from

the Cretaceous–Tertiary mass extinction. Science 282, 276–

279.

Diaz-Martinez E, Grahn Y (2007) Early Silurian glaciation

along the western margin of Gondwana (Peru, Bolivia and

northern Argentina); palaeogeographic and geodynamic set-

ting. Palaeogeography, Palaeoclimatology, Palaeoecology 245,

62–81.

Dickens GR, O’Neil JR, Rea DK, Owen RM (1995) Dissociation

of oceanic methane hydrate as a cause of the carbon isotope

excursion at the end of the Paleocene. Paleoceanography 10,

965–971.

Knoll_c21.indd 420Knoll_c21.indd 420 2/16/2012 7:49:16 PM2/16/2012 7:49:16 PM

Page 19: Fundamentals of Geobiology (Knoll/Fundamentals of Geobiology) || Geobiology of the Phanerozoic

Geobiology of the Phanerozoic 421

DiMichele WA, Pfefferkorn HW, Gastaldo RA (2001) Response

of Late Carboniferous and Early Permian plant communities

to climate change. Annual Review of Earth and Planetary Sciences 29, 461–487.

Doyle JA, Hickey LJ (1976) Pollen and leaves from the mid-

Cretaceous Potomac Group and their bearing on early angio-

sperm evolution. In: Origin and Early Evolution of the Angiosperms (ed Beck CB). Columbia University Press, New

York, pp. 139–206.

Droser ML, Bottjer DJ (1988) Trends in depth and extent of bio-

turbation in Cambrian carbonate marine environments,

Western United States. Geology 16, 233–236.

Droser ML, Bottjer DJ, Sheehan PM (1997) Evaluating the eco-

logical architecture of major events in the Phanerozoic

history of marine invertebrate life. Geology 25, 167–170.

Eagles G, Livermore R, Morris P (2006) Small basins in the

Scotia Sea; the Eocene Drake Passage gateway. Earth and Planetary Science Letters 242, 343–353.

Eldrett, JS, Harding IC, Wilson PA, Butler E, Roberts AP (2007)

Continental ice in Greenland during the Eocene and

Oligocene. Nature 446, 176–179.

Falkowski PG, Katz ME, Knoll AH, et al. (2004) The evolution

of modern eukaryotic phytoplankton. Science 305, 354–360.

Falkowski PG, Katz ME, Milligan AJ, et al. (2005) The rise of

oxygen over the past 205 million years and the evolution of

large placental mammals. Science 309, 2202–2204.

Firestone RB, West A, Kennett JP, et al. (2007) Evidence for an

extraterrestrial impact 12,900 years ago that contributed to

the megafaunal extinctions and the Younger Dryas cool-

ing. Proceedings of the National Academy of Sciences 104,

16016–16021.

Fowell SJ, Cornet B, Olsen PE (1994) Geologically rapid Late

Triassic extinctions; palynological evidence from the Newark

Supergroup. Special Paper – Geological Society of America 288,

197–206.

Freeman KH, Hayes JM (1992) Fractionation of carbon isotopes

by phytoplankton and estimates of ancient CO2 levels. Global

Biogeochemical Cycles 6,185–198.

Funakawa S, Nishi H, Moore TC, Nigrini CA (2006)

Radiolarian faunal turnover and paleoceanographic change

around Eocene/Oligocene boundary in the Central

Equatorial Pacific, ODP Leg 199, Holes 1218A, 1219A, and

1220A. Palaeogeography, Palaeoclimatology, Palaeoecology 230,

183–203.

Garrett P (1970) Phanerozoic stromatolites; noncompetitive

ecologic restriction by grazing and burrowing animals.

Science 169, 171–173.

Gentry A, Heizmann EPJ (1996) Miocene ruminants of the

central and eastern Paratethys, In: The Evolution of Western Eurasian Neogene Mammal Faunas (eds Bernor RL, Fahlbusch

V, Mittman H-W). Columbia University Press, New York,

pp. 378–395.

Graham JB, Dudley R, Aguilar NM, Gans C (1995) Implications

of the late Palaeozoic oxygen pulse for physiology and evo-

lution. Nature 375, 117–120.

Grotzinger J, Adams EW, Schroeder S (2005) Ediacaran–

Cambrian paleoecology, sedimentology and stratigraphy of

NambiaNamibia. Geological Magazine 142: 499–517.

Hagadorn JW, Bottjer DJ (1997) Wrinkle structures; microbially

mediated sedimentary structures common in subtidal silici-

clastic settings at the Proterozoic–Phanerozoic transition.

Geology 25, 1047–1050.

Hansen TA (1987) Extinction of late Eocene to Oligocene

molluscs; relationship to shelf area, temperature changes,

and impact events. Palaios 2, 69–75.

Herman AB, Spicer RA (1996) Palaeobotanical evidence for a

warm Cretaceous Arctic Ocean. Nature 380, 330–333.

Hesselbo SP, Grocke DR, Jenkyns HC, et  al. (2000) Massive

dissociation of gas hydrate during a Jurassic oceanic anoxic

event. Nature 406, 392–395.

Hickman CS (2003) Evidence for abrupt Eocene–Oligocene

molluscan faunal changes in the Pacific Northwest. In:

From  Greenhouse to Icehouse; the Marine Eocene–Oligocene Transition (eds Prothero DR, Ivany LC, Nesbitt EA). Columbia

University Press, New York, pp. 71–87.

Hill D (1972) Treatise on Invertebrate Paleontology Part E, Vol. 1 (Revised) Archaeocyatha. Geological Society of America and

University of Kansas: Boulder, Colorado and Lawrence,

Kansas), 158 pp.

Hooghiemstra H, van der Hammen T (1998) Neogene and

Quaternary development of the neotropical rain forest; the

forest refugia hypothesis, and a literature overview. Earth-Science Reviews 44, 147–183.

House MR (2002) Strength, timing, setting and cause of mid-

Paleozoic extinctions. Palaeogeography, Palaeoclimatology, Palaeoecology 181, 5–25.

Huey RB, Ward PD (2005) Hypoxia, global warming, and

terrestrial Late Permian extinctions. Science 308, 398–401.

Hulbert JRC (1993) Taxonomic evolution in North American

Neogene horses (subfamily Equinae): the rise and fall of an

adaptive radiation. Paleobiology 19, 216–234.

Isaacson PE, Díaz-Marinez E, Grader GW, Kalvoda J, Babek O,

Devuyst FX (2008) Late Devonian-earliest Mississippian gla-

ciation in Gondwanaland and its biogeographic consequences.

Palaeogeography, Palaeoclimatology, Palaeoecology 268, 126–142.

Isozaki Y (1997) Anatomy and genesis of a subduction-related

orogen: A new view of geotectonic subdivision and evolu-

tion of the Japanese Islands. Science 276, 235–238.

Isozaki Y, Kawahata H, Minoshima K (2007) The Capitanian

(Kamura) cooling event: the beginning of the Paleozoic-

Mesozoic transition. Palaeoworld 16, 16–30.

Ivany LC, van Simaeys S, Domack EW, Samson SD (2006)

Evidence for an earliest Oligocene ice sheet on the Antarctic

Peninsula. Geology 34, 377–380.

Jahren AH, Sternberg LSL (2002) Eocene meridional weather

patterns reflected in the oxygen isotopes of Arctic fossil

wood. GSA Today 12, 4–9.

Jenkyns HC (1988) The early Toarcian (Jurassic) anoxic event;

stratigraphic, sedimentary and geochemical evidence.

American Journal of Science 288, 101–151.

Jin YG, Zhang J, Shang QH (1994) Two phases of the end-

Permian mass extinction. Canadian Society of Petroleum Geologists Memoir 17, 813–822.

Joachimski MM, Buggisch W (2002) Conodont apatite δ18 O

signatures indicate climatic cooling as a trigger of the

Late Devonian mass extinction. Geology 30, 711–714.

Johnson KR, Hickey LJ (1990) Megafloral change across the

Cretaceous/Tertiary boundary in the northern Great Plains

and Rocky Mountains, USA. Special Paper - Geological Society of America 247, 433–444.

Knoll_c21.indd 421Knoll_c21.indd 421 2/16/2012 7:49:16 PM2/16/2012 7:49:16 PM

Page 20: Fundamentals of Geobiology (Knoll/Fundamentals of Geobiology) || Geobiology of the Phanerozoic

422 Fundamentals of Geobiology

Keller G (1983) Paleoclimatic analyses of middle Eocene

through Oligocene planktic foraminiferal faunas.

Palaeogeography, Palaeoclimatology, Palaeoecology 43, 73–94.

Kennett, DJ, Kennett JP, West A, et al. (2009) Shock-synthesized

hexagonal diamonds in Younger Dryas boundary sediments.

Proceedings of the National Academy of Sciences 106, 12623–12628.

Kennett JP, Stott LD (1991) Abrupt deep-sea warming, palae-

oceanographic changes and benthic extinctions at the end of

the Palaeocene. Nature 353, 225–229.

Kennett JP, Cannariato KG, Hendy IL, Behl R (2003) Methane Hydrates in Quaternary Climate Change; the Clathrate Gun Hypothesis. American Geophysical Union, Washington, DC.

Kennett JP, Houtz RE, Andrews PB, et al. (1975) Cenozoic pale-

oceanography in the Southwest Pacific Ocean, Antarctic gla-

ciation, and the development of the Circum-Antarctic

Current. Initial Reports of the Deep Sea Drilling Project 29, 1155–1169.

Knight KB, Nomade S, Renne PR, Marzoli A, Bertrand H,

Youbi N (2004) The Central Atlantic magmatic province at

the Triassic-Jurassic boundary; paleomagnetic and 40Ar/39Ar

evidence from Morocco for brief, episodic volcanism. Earth and Planetary Science Letters 228, 143–160.

Knoll AH, Bambach RK, Payne JL, Pruss S, Fischer WW (2007a)

Paleophysiology and end-Permian mass extinction. Earth and Planetary Science Letters 256, 295–313.

Knoll AH, Summons RE, Walbauer JR, Zumberge JE (2007b)

The Geological Succession of Primary Producers in the

Oceans. In Evolution of Primary Producers in the Sea (eds

Falkowski PG, Knoll AH). Elsevier Acacemic Press,

Amsterdam, pp. 133–163.

Kobashi T, Grossman EL, Yancey TE, Dockery DT, III (2001)

Reevaluation of conflicting Eocene tropical temperature esti-

mates; molluskan oxygen isotope evidence for warm low

latitudes. Geology 29, 983–986.

Kobashi T, Grossman EL, Dockery DT, III, Ivany LC (2004)

Water mass stability reconstructions from greenhouse

(Eocene) to icehouse (Oligocene) for the northern Gulf Coast

continental shelf (USA). Paleoceanography 19, 16.

Kump LR, Pavlov A, Arthur MA (2005) Massive release of

hydrogen sulfide to the surface ocean and atmosphere

during intervals of oceanic anoxia. Geology 33, 397–400.

Kuypers MMM, van Breugel Y, Schouten S, Erba E, Sinninghe

Damsté JS (2004) N2-fixing cyanobacteria supplied nutrient

N for Cretaceous oceanic anoxic events. Geology 32, 853–856.

Labandeira CC, Phillips TL, Norton RA (1997) Oribatid mites

and the decomposition of plant tissues in Paleozoic coal-

swamp forests. Palaios 12, 319–353.

Larson RL (1991) Latest pulse of Earth; evidence for a Mid-

Cretaceous super plume. Geology 19, 547–550.

Leckie RM, Bralower TJ, Cashman R (2002) Oceanic anoxic

events and plankton evolution; biotic response to tectonic

forcing during the Mid-Cretaceous. Paleoceanography 17,

no. 3, 29pp.

Lisse, CM, Kraemer KE, Nuth JA, Li A, Josiak, D (2007)

Comparison of the composition of the Tempel 1 ejecta to the

dust in Comet C/Hale-Bopp 1995 O1 and YSO HD 100546.

Icarus 191, 223–240.

Loydell DK (2007) Early Silurian positive δ13 C excursions and

their relationship to glaciations, sea-level changes and

extinction events. Geological Journal 42, 531–546.

Magioncalda R, Dupuis C, Smith T, Steurbaut E, Gingerich PD

(2004) Paleocene–Eocene carbon isotope excursion in organic

carbon and pedogenic carbonate; direct comparison in a con-

tinental stratigraphic section. Geology 32, 553–556.

Maliva RG, Knoll AH, Siever R (1989) Secular change in chert

distribution: a reflection of evolving biological participation

in the silica cycle. Palaios 4, 519–532.

Marubini F, Thake B (1999) Bicarbonate addition promotes

coral growth. Limnology and Oceanography 44, 716–720.

Marzoli A, Renne PR, Piccirillo EM, Ernesto M, Bellieni G, De

Min A (1999) Extensive 200-million-year-old continental

flood basalts of the Central Atlantic Magmatic Province.

Science 284, 616–618.

McElwain JC, Beerling DJ, Woodward FI (1999) Fossil plants

and global warming at the Triassic-Jurassic boundary. Science

285, 1386–1390.

Melosh HJ, Schneider NM, Zahnle KJ, Latham D (1990) Ignition

of global wildfires at the Cretaceous/Tertiary boundary.

Nature 343, 251–254.

Meng J, McKenna MC (1998) Faunal turnovers of Paleogene

mammals from the Mongolian Plateau. Nature 394, 364–367.

Montanari A, Hay RL, Alvarez W, et al. (1983) Spheroids at the

Cretaceous–Tertiary boundary are altered impact droplets of

basaltic composition. Geology 11, 668–671.

Olsen PE, Kent DV, Sues HD, et al. (2002) Ascent of dinosaurs

linked to an iridium anomaly at the Triassic–Jurassic bound-

ary. Science 296, 1305–1307.

Pagani M, Zachos JC, Freeman KH, Tipple B, Bohaty S (2005)

Marked decline in atmospheric carbon dioxide concentra-

tions during the Paleogene. Science 309, 600–603.

Palmer AR (1984) The biomere problem; evolution of an idea.

Journal of Paleontology 58, 599–611.

Palmer AR (1998) A proposed nomenclature for stages and

series for the Cambrian of Laurentia. Canadian Journal of Earth Sciences 35, 323–328.

Paquay FS, Goderis S, Ravizza G. et al. (2009) Absence of geo-

chemical evidence for an impact event at the Bolling–

Allerod/Younger Dryas transition. Proceedings of the National Academy of Sciences 106, 21505–21510.

Parrish JT, Spicer RA (1988) Late Cretaceous terrestrial vegeta-

tion; a near-polar temperature curve. Geology 16, 22–25.

Payne JL, Lehrmann DJ, Wei J, Orchard MJ, Schrag DP, Knoll

AH (2004) Large perturbations of the carbon cycle during

recovery from the end-Permian extinction. Science 305,

506–509.

Perfetta PJ, Shelton KV, Stitt JH (1999) Carbon isotope evidence

for deep-water invasion at the marjumiid-pterocephaliid

biomere boundary, Black Hills, USA: A common origin for

biotic crises on Late Cambrian shelves. Geology 27, 403–406.

Pochon X, Montoya-Burgos JI, Stadelmann B, Pawlowski J

(2006) Molecular phylogeny, evolutionary rates, and diver-

gence timing of the symbiotic dinoflagellate genus

Symbiodinium. Molecular Phylogenetics and Evolution 38,

20–30.

Poore RZ, Berggren WA (1975) Late Cenozoic planktonic

foraminiferal biostratigraphy and paleoclimatology of

Hatton-Rockall Basin; DSDP Site 116. Journal of Foraminiferal Research 5, 270–293.

Pope KO, Baines KH, Ocampo AC, Ivanov BA (1994) Impact

winter and the Cretaceous/Tertiary extinctions; results of a

Knoll_c21.indd 422Knoll_c21.indd 422 2/16/2012 7:49:16 PM2/16/2012 7:49:16 PM

Page 21: Fundamentals of Geobiology (Knoll/Fundamentals of Geobiology) || Geobiology of the Phanerozoic

Geobiology of the Phanerozoic 423

Chicxulub asteroid impact model. Earth and Planetary Science Letters 128, 719–725.

Porter SM (2007) Seawater chemistry and early carbonate

biomineralization. Science 316, 302.

Powell MG (2005) Climatic basis for sluggish macroevolution

during the late Paleozoic ice age. Geology 33, 381–384.

Puceat E, Lecuyer C, Donnadieu Y, et al. (2007) Fish tooth δ18 O

revising Late Cretaceous meridional upper ocean water

temperature gradients. Geology 35, 107–110.

Raffi S, Stanley SM, Marasti R (1985) Biogeographic patterns and

Plio-Pleistocene extinction of Bivalvia in the Mediterranean

and southern North Sea. Paleobiology 11, 368–388.

Rees, DL (2002) Land-plant diversity and the end-Permian

mass extinction. Geology 30: 827–830

Reimers CE, Jahnke RA, McCorkle DC (1992) Carbon fluxes

and burial rates over the continental slope and rise off

Central California with implications for the global carbon

cycle. Global Biogeochemical Cycles 6,199–224.

Renegar DA, Riegl BM (2005) Effect of nutrient enrichment and

elevated CO2 partial pressure on growth rate of Atlantic scle-

ractinian coral Acropora cervicornis. Marine Ecology Progress Series 293, 69–76.

Retallack GJ (1997) Early forest soils and their role in Devonian

global change. Science 276, 583–585.

Retallack GJ, Metzger CA, Greaver T, Jahren AH, Smith RMH,

Sheldon ND (2006) Middle-Late Permian mass extinction on

land. Geological Society of America Bulletin 118, 1398–1411.

Ries JB, Stanley SM, Hardie LA (2006) Scleractinian corals pro-

duce calcite, and grow more slowly, in artificial Cretaceous

seawater. Geology 34, 525–528.

Royer DL (2003) Estimating latest Cretaceous and Tertiary

atmospheric CO2 from stromatal indices. Geological Society of

America Special Paper 369, 79–93.

Royer DL, Berner RA, Beerling DJ (2001) Phanerozoic atmos-

pheric CO2 change: evaluating geochemical and paleobio-

logical approaches. Earth-Science Reviews 54, 349–392.

Saito T, Yamanoi T, Kaiho K (1986) End-Cretaceous devastation

of terrestrial flora in the boreal Far East. Nature 323, 253–255.

Saltzman MR (2001) Silurian δ13 C stratigraphy: A view from

North America. Geology, 29, 671–674.

Saltzman MR (2002) Carbon isotope (δ13 C) stratigraphy across

the Silurian–Devonian transition in North America: evi-

dence for a perturbation of the global carbon cycle.

Palaeogeography, Palaeoclimatology, Palaeoecology 187, 83–100.

Saltzman MR, Young SA (2005) Long-lived glaciation in the

Late Ordovician? Isotopic and sequence-stratigraphic evi-

dence from western Laurentia. Geology 33, 109–112.

Scher HD, Martin EE (2006) Timing and climatic consequences

of the opening of Drake Passage. Science 312, 428–430.

Schouten S, Hopmans EC, Forster A, van Breugel Y, Kuypers

MMM, Sinninghe Damsté JS (2003) Extremely high sea-

surface temperatures at low latitudes during the Middle

Cretaceous as revealed by archaeal membrane lipids. Geology

31, 1069–1072.

Sepkoski JJ (1982) Flat-pebble conglomerates, storm deposits,

and the Cambrian bottom fauna. In: Cyclic and Event Stratification (eds Einsele G, Seilacher A). Springer-Verlag,

Berlin, pp. 371–385.

Sheehan PM (2001) The Late Ordovician mass extinction.

Annual Review of Earth and Planetary Sciences 29, 331–364.

Soreghan GS, Soreghan MJ, Poulsen CJ, et al. (2008) Anomalous

cold in the Pangaean tropics. Geology 36, 659–662.

Stanley SM (1986) Anatomy of a regional mass extinction: Plio-

Pleistocene decimation of the Western Atlantic bivalve

fauna. Palaios 1, 17–36.

Stanley SM (1990) Adaptive radiation and macroevolution.

Systematics Association Special Volume 42, 1–16.

Stanley SM (1992) An ecological theory for the origin of Homo. Paleobiology 18, 237–257.

Stanley SM (1995) New horizons for paleontology, with two

examples; the rise and fall of the Cretaceous Supertethys and

the cause of the modern ice age. Journal of Paleontology 69,

999–1007.

Stanley SM (2009a) Earth System History. W.H.Freeman and

Company, New York.

Stanley SM (2009b) Evidence from ammonoids and conodonts

for multiple Early Triassic mass extinctions. Proceedings of the National Academy of Sciences 106, 15256–15259.

Stanley SM (2010) Relation of Phanerozoic stable isotope

excursions to climate, bacterial metabolism, and major

extinctions. Proceedings of the National Academy of Sciences

107, 19185–19189.

Stanley SM, Yang X (1994) A double mass extinction at the end

of the Paleozoic Era. Science 266, 1340–1344.

Stanley SM, Hardie LA (1998) Secular oscillations in carbonate

mineralogy of reef-building and sediment-producing organ-

isms driven by tectonically forced shifts in seawater chemis-

try. Palaeogeography, Palaeoclimatology, Palaeoecology 144, 3–19.

Stanley SM, Powell MG (2003) Depressed rates of origination

and extinction during the late Paleozoic ice age; a new state

for the global marine ecosystem. Geology 31, 877–880.

Stanley SM, Ries JB, Hardie LA (2005) Seawater chemistry, coc-

colithophore population growth, and the origin of Cretaceous

chalk. Geology 33, 593–596.

Stickley CE, Brinkhuis H, Schellenberg SA, et al. (2004) Timing

and nature of the deepening of the Tasmanian Gateway.

Paleoceanography 19, 18.

Stitt JH (1975) Adaptive radiation, trilobite paleoecology, and

extinction, ptychaspidid biomere, late Cambrian of

Oklahoma. Fossils and Strata 4, 381–390.

Stromberg CAE (2004) Using phytolith assemblages to recon-

struct the origin and spread of grass-dominated habitats in

the Great Plains of North America during the late Eocene to

early Miocene. Palaeogeography, Palaeoclimatology, Palaeoecology

207, 239–275.

Swisher CC, Grajales-Nishimura JM, Montanari A, et al. (1992)

Coeval 40Ar/39Ar ages of 65.0 million years ago from

Chicxulub Crater melt rock and Cretaceous–Tertiary bound-

ary tektites. Science 257, 954–958.

Tanner LH, Lucas SG, Chapman MG (2004) Assessing the

record and causes of Late Triassic extinctions. Earth-Science Reviews 65, 103–139.

Tschudy RH, Tschudy BD (1986) Extinction and survival of

plant life following the Cretaceous/Tertiary boundary event,

Western Interior, North America. Geology 14, 667–670.

Vrba ES (1985) African Bovidae: evolutionary events since the

Miocene. South African Journal of Science 81, 263–266.

Webb SD (1977) A history of the savanna vertebrates in the

New World; Part I, North America. Annual Review of Ecology and Systematics 8, 355–380.

Knoll_c21.indd 423Knoll_c21.indd 423 2/16/2012 7:49:16 PM2/16/2012 7:49:16 PM

Page 22: Fundamentals of Geobiology (Knoll/Fundamentals of Geobiology) || Geobiology of the Phanerozoic

424 Fundamentals of Geobiology

Webb SD (1984) Ten million years of mammal extinctions in

North America. In Quaternary Extinctions: A Prehistoric Revolution (eds Martin PS, Klein RG). University of Arizona

Press, Tucson, pp. 189–210.

Whiteside JH, Olsen PE, Kent DV, Fowell SJ, Et-Touhami M

(2007) Synchrony between the Central Atlantic magmatic

province and the Triassic-Jurassic mass-extinction event?

Palaeogeography, Palaeoclimatology, Palaeoecology 244, 345–367.

Wing SL, Harrington GJ, Smith FA, Bloch JI, Boyer DM,

Freeman KH (2005) Transient floral change and rapid global

warming at the Paleocene–Eocene boundary. Science 310,

993–996.

Wing SL, Hickey LJ, Swisher CC (1993) Implications of an

exceptional fossil flora for Late Cretaceous vegetation.

Nature 363, 342–344.

Wolfe JA (1971) Tertiary climatic fluctuations and methods of

analysis of Tertiary floras. Palaeogeography, Palaeoclimatology, Palaeoecology 9, 27–57.

Zachos JC, Roehl U, Schellenberg SA, et al. (2005) Rapid acidi-

fication of the ocean during the Paleocene–Eocene thermal

maximum. Science 308, 1611–1615.

Zachos JC, Wara MW, Bohaty S, et al. (2003) A transient rise in

tropical sea surface temperature during the Paleocene-

Eocene thermal maximum. Science 302, 1551–1554.

Zhu M-Y, Zhang, J-M,Li G-X, Yang, A-H (2004) Evolution of C

isotopes in the Cambrian of China: implications for

Cambrian subdivision and trilobite mass extinctions.

Geobios 37, 287–301.

Zhuravlev AY, Wood R (2008) Eve of biomineralization:

Controls on skeletal mineralogy. Geology 36, 923–926.

Knoll_c21.indd 424Knoll_c21.indd 424 2/16/2012 7:49:16 PM2/16/2012 7:49:16 PM