fundamentals of geobiology (knoll/fundamentals of geobiology) || geobiology of the phanerozoic
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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
Knoll_c21.indd 409Knoll_c21.indd 409 2/16/2012 7:49:13 PM2/16/2012 7:49:13 PM
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).
Knoll_c21.indd 410Knoll_c21.indd 410 2/16/2012 7:49:13 PM2/16/2012 7:49:13 PM
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).
Knoll_c21.indd 411Knoll_c21.indd 411 2/16/2012 7:49:13 PM2/16/2012 7:49:13 PM
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).
Knoll_c21.indd 412Knoll_c21.indd 412 2/16/2012 7:49:14 PM2/16/2012 7:49:14 PM
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).
Knoll_c21.indd 413Knoll_c21.indd 413 2/16/2012 7:49:14 PM2/16/2012 7:49:14 PM
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
Knoll_c21.indd 415Knoll_c21.indd 415 2/16/2012 7:49:14 PM2/16/2012 7:49:14 PM
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).
Knoll_c21.indd 416Knoll_c21.indd 416 2/16/2012 7:49:14 PM2/16/2012 7:49:14 PM
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).
Knoll_c21.indd 417Knoll_c21.indd 417 2/16/2012 7:49:14 PM2/16/2012 7:49:14 PM
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
Knoll_c21.indd 418Knoll_c21.indd 418 2/16/2012 7:49:15 PM2/16/2012 7:49:15 PM
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).
Knoll_c21.indd 419Knoll_c21.indd 419 2/16/2012 7:49:15 PM2/16/2012 7:49:15 PM
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.
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