Chapter 8

Clocks in rocks
Reconstructing Geologic History from the Stratigraphic Record
Geologists speak carefully about time. To them, dating refers to measuring the absolute age of an
event in the geologic record. The number of years elapsed from that event until now. Before the
twentieth century no one knew much about absolute ages. They could say that fish bones were first
deposited in marine sediments before mammal bones first appeared in sediments on land but they
couldn’t tell how many years ago. The first geologic observations pertaining to the question of deep
time came in the mid seventeen century from the study of fossils. A fossil is an artifact of life
preserved in the geologic record. However few people living in seventeen century Europe would
have understood this definition. Most thought that the seashells and other lifelike forms they found
in rocks dated from earths beginnings about 6000 years earlier, or grew there spontaneously.

In 1667 the Danish scientists Nicolaus Steno who was working for the royal court in Florence Italy,
demonstrated that the peculiar tongue stones found in certain Mediterranean sedimentary rocks
were essentially identical to the teeth of modern sharks. He concluded that tongue stones really are
ancient shark teeth preserved in the rocks and more generally that fossils were the remains of
ancient life deposited with sediments.

Principles of Stratigraphy
Geologists still use the principles set forth by Steno to interpret sedimentary strata. Two of his basic
rules are so simple they seem obvious to us today:

1. The principle of original horizontality states that sediments are deposited under the influence of
gravity as nearly horizontal beds. Observations in a wide variety of sedimentary environments
support this principle. If we find folded or faulted strata, we know that the beds were deformed by
tectonic forces after the sediments were deposited.
 2. The principle of superposition states that each layer of an unreformed sedimentary sequence is
younger than the one beneath it and older than the one above it. A new layer cannot be deposited
beneath an existing layer. Thus, strata can be vertically ordered in time from the lowest (oldest) bed
to the uppermost (youngest) bed (Figure 8.4). A chronologically ordered set of strata is known as a
stratigraphic succession.

We can apply Steno's principles in the field to deter-mine whether one sedimentary formation is
older than another. Then, by piecing together the formations exposed in different outcrops, we can
sort them into chronological order and thus construct the stratigraphic succession of a region—at
least in principle. In practice, there were two problems with this strategy. First, geologists almost
always found gaps in a region's stratigraphic succession, indicating time intervals that had gone
entirely unrecorded. Some of these intervals were short, such as periods of drought between floods;
others lasted for millions of years.

Example, periods of regional tectonic uplift when thick sequences of sedimentary rocks were
removed by erosion. Second, it was difficult to determine the relative ages of two formations that
were widely separated in space; stratigraphy alone couldn't determine whether a sequence of
mudstones in, say, Tuscany was older, younger, or the same age as a similar sequence in England. It
was necessary to expand Steno's ideas about the biological origin of fossils to solve these problems.

Fossils as Recorders of Geologic Time

In 1793, William Smith, a surveyor working on the construction of canals in southern England,
recognized that fossils could help geologists determine the relative ages of sedimentary rocks. Smith
was fascinated by the variety of fossils he collected from the strata exposed along canal cuts. He
observed that different layers contained different sets of fossils, and he was able to tell one layer
from another by the characteristic fossils in each. He established a general order for the sequence of
fossil assemblages and strata, from lowest (oldest) to uppermost (youngest). Regardless of its
location, Smith could predict the stratigraphic position of any particular layer or formation in any
outcrop in southern England based on its fossil assemblages. This stratigraphic ordering of the fossils
of animal species (fauna) produces a sequence known as a faunal succession.

Smith's principle of faunal succession states that the sedimentary strata in an outcrop contain fossils
in a definite sequence. The same sequence can be found in out-crops at other locations, so that
strata in one location can be matched to strata in another location.
 Using faunal successions, Smith was able to identify formations of the same age in different
outcrops. By noting the vertical order in which the formations were found in each place, he compiled
a composite stratigraphic succession for the entire region. His compilation showed how the
complete succession would have looked if the formations at different levels in all the various
outcrops could have been brought together at a single spot. Figure 8.5 shows such a composite
stratigraphic succession for two outcrops.

Smith kept track of his work by mapping outcrops using colors assigned to specific formations, thus
inventing the geologic map (see Figure 7.4). In 1815, he summarized his lifelong research by
publishing his General Map of Strata in England and Wales, a hand-colored masterpiece 8 feet tall
and 6 feet wide—the first geologic map of an entire country. The original still hangs in the offices of
the Geological Society of London.

The geologists who followed in Steno's and Smith's footsteps described and catalogued hundreds of
fossils and their relationships to modern organisms, establishing the new science of paleontology:
the historical study of ancient life-forms. The most common fossils they found were the shells of
invertebrate animals. Some were similar to clams, oysters, and other living shellfish; others
represented strange species with no living examples, such as the trilobites shown in the chapter
opening photo. Less common were the bones of vertebrates, such as mammals, birds, and the huge
extinct reptiles they called dinosaurs. Plant fossils were found to be abundant in some rocks,
particularly in coal beds, where leaves, twigs, branches, and even whole tree trunks could be
recognized. Fossils were not found in intrusive igneous rocks—no surprise, since any biological
material would have been destroyed when the rocks melted—nor in high-grade metamorphic
rocks—where any remains of organisms would have been distorted be-yond recognition.

By the beginning of the nineteenth century, paleontology had become the single most important
source of information about geologic history. The systematic study

200 CHAPTER 8 Clocks in Rocks: Timing the Geologic Record of fossils affected science far beyond
geology, however. Charles Darwin studied paleontology as a young scientist, and he collected many
unusual fossils on his famous voyage aboard the Beagle (1831-1836). During this world-circling tour,
he also studied many unfamiliar animal and plant species in their native habitats. Darwin pondered
what he had seen until 1859, when he proposed his theory of evolution by natural selection. His
theory revolutionized the science of biology and provided a sound theoretical framework for
paleontology: if organisms evolve progressively over time, then the fossils in each sedimentary bed
must represent the organisms living when that bed was deposited.

Unconformities: Gaps in the Geologic Record

In compiling the stratigraphic succession of a region, geologists often find places in the geologic
record where a formation is missing. Either no rock was ever deposited, or it was eroded away
before the next strata were laid down. The surface between two beds that were laid down with a
time gap between them—the boundary representing the missing time—is called an unconformity
(Figure 8.6). A series of beds bounded above and below by unconformities is referred to as a
sedimentary sequence.

An unconformity, like a sedimentary sequence, represents the passage of time. An unconformity

may imply that tectonic forces raised the rock above sea level, where erosion removed some rock
layers. Alternatively, the unconformity may have been produced by the erosion of newly exposed
rock as sea level fell. As we will see in Chapter 21, global sea level can be lowered by hundreds of
meters during ice ages, when water is withdrawn from the oceans to form continental ice sheets.
Unconformities are classified according to the relation-ships between the layers above and below
them:

 A disconformity is an unconformity in which an upper sedimentary sequence overlies an

erosional surface developed on an underformed, still-horizontal lower sedimentary
sequence (see Figure 8.6). Disconformities are often created when sea level drops or during
broad tectonic uplifts.
 A nonconformity is an unconformity in which the upper sedimentary beds overlie
metamorphic or igneous rocks (see Earth Issues 8.1, pages 204-205, for an example).
 An angular unconformity is an unconformity in which the upper beds overlie lower beds
that have been folded by tectonic processes and then eroded to a more or less even plane.
In an angular unconformity, the two sequences have bedding planes that are not parallel.
Figure 8.7 depicts a dramatic angular unconformity found near the bottom of the Grand
Canyon. The formation of an angular unconformity by tectonic processes is illustrated in
Figure 8.8.
Cross-Cutting Relationships
Other disturbances of the layering of sedimentary strata also provide clues for determining the
relative ages of rocks. Recall that dikes can cut through sedimentary beds; sills can be intruded
parallel to bedding planes (see Chapter 4); and faults can displace bedding planes, dikes, and sills as
they shift blocks of rock (see Chapter 7). These cross-cutting relationships can be used to establish
the relative ages of igneous intrusions or faults within the stratigraphic suc-cession. Because the
deformation or intrusion events must have taken place after the affected sedimentary beds were
deposited, those structures must be younger than the rocks they cut (Figure 8.9). If the intrusions or
fault displace-ments are eroded and planed off at an unconformity and then overlaid by younger
sedimentary beds, we know that those structures are older than the younger strata.
Geologists can combine field observations of cross-cutting relationships, unconformities, and
stratigraphic successions to decipher the history of geologically complicated regions (Figure 8.10).
Earth Issues 8.1 gives a more detailed example of how geologists work backward in time to
determine the relative ages of the rocks in a region.
The Geologic Time Scale: Relative Ages
Early in the nineteenth century, geologists began to apply Steno's and Smith's stratigraphic
principles to outcrops all over the world. The same distinctive fossils were discovered in similar
formations on many continents. Moreover, faunal successions from different continents often
displayed the same changes in fossil assemblages. By matching up faunal successions and using
cross-cutting relationships, geologists could determine the relative ages of rock formations on a
global basis. By the end of the century, they had pieced together a worldwide history of geologic
events—a geologic time scale.

Intervals of Geologic Time

The geologic time scale divides Earth's history into intervals marked by distinct sets of fossils, and it
places the bound-aries of those intervals at times when those sets of fossils changed abruptly (Figure
8.11). The basic divisions of this time scale are the eras: the Paleozoic (from the Greek paleo,
meaning "old," and zoi , meaning "life"), the Mesozoic ("middle life"), and the Cenozoic ("new life").
The eras are subdivided into periods, usually named for the locality in which the formations
representing them were first or best described, or for some distinguishing characteristic of the
formations. The Jurassic period, for example, is named for the Jura mountain range of France and
Switzerland, and the Carboniferous period is named for the coal-bearing sedimentary rocks of
Europe and North America. The Paleogene and Neogene periods of the Cenozoic are two exceptions:
these Greek names mean "old origin" and "new origin, “respectively. Some periods are further
subdivided into epochs, such as the Miocene, Pliocene, and Pleistocene epochs of the Neogene
period (see Figure 8.11). Today we are living in the Holocene ("completely new") epoch of the
Neogene period in the Cenozoic era.

Interval Boundaries Mark Mass Extinctions

Many of the major boundaries in the geologic time scale represent mass extinctions: short intervals
during which a large proportion of the species living at the time simply disappeared from the fossil
record, followed by the blossoming of many new species. These abrupt changes in faunal
successions were a great mystery to the geologists who discovered them. Darwin's theory of
evolution explained how new species could evolve, but what had caused the mass extinctions?

In some cases, we think we know. The mass extinction at the end of the Cretaceous period, which
killed off 75 percent of the living species, including all the dinosaurs, was almost certainly the result
of a large meteorite impact that darkened and poisoned the atmosphere and plunged Earth's
climate into many years of bitter cold. This disaster marks the end of the Mesozoic era and the
beginning of the Cenozoic. In other cases, we are still not sure. The largest mass extinction, at the
end of the Permian period, which defines the Paleozoic-Mesozoic boundary, eliminated nearly 95
percent of all living species, but the cause of this event is still debated. The extreme events that
separate intervals of geologic time are the subject of very active research, as we will see in Chapter
11.

Ages of Petroleum Source Rocks

Oil and natural gas come from organic matter that was buried in sedimentary rock formations at
some time in the geologic past. The relative ages of these "petroleum source rocks" provide
important clues about where to look for new oil and gas resources. Global surveys have shown that
very little petroleum has come from Precambrian rocks, which makes sense, because the primitive
organ-isms that existed before the Cambrian period generated little organic matter.

Petroleum source rocks were deposited during all three of the geologic eras following the Cambrian,
although certain periods of geologic time have produced much more of this resource than others
(Figure 8.12). The clear winners are the Jurassic and Cretaceous periods of the Mesozoic era, which
together have accounted for almost 60% of the world's petroleum production. Sedimentary
formations of Jurassic and Cretaceous age were the source rocks for the great oil fields of the Middle
East, the Gulf of Mexico, Venezuela, and the North Slope of Alaska.

If you examine Figure 2.16, you can see that, during these periods of geologic time, the
supercontinent of Pangaea was breaking up into the modern continents. This tectonic activ-ity
formed many marine sedimentary basins and increased the rate at which sediments were deposited
into these basins. During the Jurassic and Cretaceous periods, which comprise the Age of Dinosaurs,
marine life was abundant, providing much of the organic matter that was buried in the sediments.
This carbon-rich material has since been "cooked" and trans-ported into the oil reservoirs, where we
find it today.
Measuring Absolute Time with Isotopic Clocks
The geologic time scale based on stratigraphy and faunal successions is a relative time scale. It tells
us whether one formation or fossil assemblage is older than another, but not how long the eras,
periods, and epochs were in actual years.

Estimates of how long it takes mountains to erode and sediments to accumulate suggested that
most geologic periods had lasted for millions of years, but nineteenth-century geologists did not
know whether the duration of any specific period was 10 million years, 100 million years, or even
longer.

They did know that the geologic time scale was incomplete. The earliest period of geologic history
recorded by faunal successions was the Cambrian, when animal life, in the form of shelly fossils,
suddenly appeared in the geologic record. Many rock formations were clearly older, however,
because they occurred below Cambrian rocks in strati-graphic successions. But these formations
contained no recognizable fossils, so there was no way to determine their relative ages. All such
rocks were lumped into the general category Precambrian. What fraction of Earth's history was
locked up in these cryptic rocks? How old was the oldest Precambrian rock? How old was Earth
itself?
These questions sparked a huge debate in the latter half of the nineteenth century. Physicists and
astronomers argued for a maximum age of less than 100 million years, but most geologists regarded
this age as much too young, even though they had no precise data to back them up.

Discovery of Radioactivity
In 1896, a major advance in physics paved the way for reliable and accurate measurements of
absolute ages. Henri Becquerel, a French physicist, discovered radioactivity in uranium. Within three
years, the French chemist Marie Sklodowska-Curie discovered and isolated a new and highly
radioactive element, radium.

In 1905, the physicist Ernest Rutherford suggested that the absolute age of a rock could be
determined by measuring the decay of radioactive elements found in it. He calculated the age of one
rock from measurements of its uranium content. This was the start of isotopic dating, the use of
naturally occurring radioactive elements to determine the ages of rocks. Isotopic dating methods
were refined over the next few years as more radioactive elements were discovered and the
processes of radioactive decay became better understood. Within a decade of Rutherford's first
attempt, geologists were able to show that some Precambrian rocks were billions of years old.

In 1956, the geochemist Clair Patterson measured the decay of uranium in meteorites and
terrestrial rocks to determine that the solar system—and, by implication, Earth—was formed 4.56
billion years ago. That age has been modified by less than 10 million years since Patterson's original
measurement, so we might say that he completed the discovery of geologic time.

Radioactive Isotopes: The Clocks in Rocks

How do geologists use radioactivity to determine the age of a rock? Recall that the nucleus of an
atom consists of protons and neutrons. For a given element, the number of protons is constant, but
the number of neutrons can vary among different isotopes of the same element (see Chapter 3).
Most isotopes are stable, but the nucleus of a radioactive isotope can spontaneously disintegrate, or
decay, emitting particles and transforming the atom into an atom of a dif-ferent element. We call
the original atom the parent and the product of decay its daughter.

One useful element for isotopic dating is rubidium, which has 37 protons and two naturally
occurring isotopes: rubidium-85, which has 48 neutrons and is stable, and rubidium-87, which has 50
neutrons and is radioactive. A neutron in the nucleus of a rubidium-87 atom can spon-taneously
emit an electron, thus changing into a proton, which remains in the nucleus. The parent rubidium
atom thus forms a daughter strontium-87 atom, with 38 protons and 49 neutrons (Figure 8.13).

A parent isotope decays into a daughter isotope at a constant rate. The rate of radioactive decay is
mea-sured by the isotope's half-life: the time required for one-half of the original number of parent
atoms to be transformed into daughter atoms. At the end of the first half-life, the number of parent
atoms has decreased a factor of two; at the end of the second half-life, by factor of four; at the end
of the third half-life, by a fact of eight, and so forth. As the parent decays, the amount of the
daughter isotope increases, preserving the tot number of atoms (Figure 8.14). The half-lives of radii
active elements commonly used for isotopic dating a given in Table 8.1.
Radioactive isotopes make good clocks because the half-lives do not vary with the changes in
temperature pressure, chemical environment, or other factors that c. accompany geologic processes
on Earth or other plane So when atoms of a radioactive isotope are created an where in the
universe, they start to like a ticking do steadily transforming from one type of atom to another a
fixed rate.

We can measure the ratio of parent to daughter ato in a rock sample with a mass spectrometer—a
precise sensitive instrument that can detect even minute quantities of isotopes—and determine
how much of the daughter has been produced from the parent. Knowing the half-life, we can then
calculate the time elapsed since the isotopic clock began to tick. The isotopic age of a rock
corresponds to the time since the isotopic clock was "reset" when the isotopes were locked into the
minerals of the rock. This "locking" usually occurs when a mineral crystallizes from magma or re-
crystallizes during metamorphism. During crystallization, however, the number of daughter atoms in
a mineral is not necessarily reset to zero, so the initial number of daughter atoms must be taken into
account when calculating isotopic age (see the Practicing Geology exercise at the end of the
chapter). Many other complications make isotopic dating a tricky business. A mineral can lose
daughter isotopes by weathering or be contaminated by fluids circulating in the rock.
Metamorphism of igneous rocks can reset the isotopic age of minerals in those rocks to a date much
later than their crystallization age.

Isotopic Dating Methods

Isotopic dating is possible only if a measurable number of parent and daughter atoms remain in the
sample being dated. For example, if a rock is very old and the decay rate of an isotope is fast, almost
all the parent atoms will already have been transformed. In that case, we could determine that the
isotopic clock had run down, but we would not be able to say when. Thus, isotopes that decay slowly
over billions of years, such as rubidium-87, are most useful in measuring the ages of older rocks,
whereas those that de-cay rapidly, such as carbon-14, can only be used to date younger rocks (see
Table 8.1).

Carbon-14, which has a half-life of about 5700 years, is especially useful for dating bone, shell, wood,
and other organic materials in sediments less than a few tens of thou-sands of years old. Carbon is
an essential element in the living cells of all organisms. As green plants grow, they con-tinuously
incorporate carbon into their tissues from carbon dioxide in the atmosphere. When a plant dies,
however, it stops absorbing carbon dioxide. At the moment of death, the ratio of carbon-14 to the
stable isotope carbon-12 in the plant is identical to that in the atmosphere. Thereafter, the ratio
decreases as the carbon-14 in the dead tissue decays. Nitrogen-14, the daughter isotope of carbon-
14, is a gas and thus leaks from the material, so it cannot be measured to determine the time that
has elapsed since the plant died. We can, however, estimate the absolute age of the plant material
by comparing the ratio of carbon-14 left in the plant material with the ratio in the atmosphere at the
time the plant died. The latter ratio can be estimated from carbon-14 ages calibrated using other
measures of absolute time, such as dendrochronology (counting tree rings).

One of the most precise dating methods for old rocks is based on the decay of two related isotopes:
the decay of uranium-238 to lead-206 and the decay of uranium-235 to lead-207. Isotopes of the
same element behave similarly in the chemical reactions that alter rocks because the chemistry of an
element depends mainly on its atomic number, not its atomic mass. The two uranium isotopes have
different half-lives, however, so together they provide a consistency check that helps geologists
compensate for the problems of weathering, contamination, and metamorphism discussed above.
The lead isotopes from single crystals of zircon—zirconium silicate, a crustal mineral with a relatively
high concentration of uranium—can be used to date the oldest rocks on Earth with an uncertainty of
less than 1 percent. These formations turn out to be more than 4 billion years old.

The Geologic Time Scale: Absolute Ages

Armed with isotopic dating techniques, geologists of the twentieth century were able to nail down
the absolute ages of the key events on which their predecessors had based the geologic time scale.
More important, they were able to explore the early history of the planet recorded in Precambrian
rocks. Figure 8.15 presents the results of this century-long effort. The assignment of absolute ages to
the geologic time scale revealed great differences in the lengths of the geo-logic periods. The
Cretaceous period (spanning 80 mil-lion years) turned out to be more than three times longer than
the Neogene period (only 23 million years), and the Paleozoic era (291 million years) was found to
be longer than the Mesozoic and Cenozoic eras combined. The big-gest surprise was the
Precambrian, which had a duration of over 4000 million years—almost nine-tenths of Earth's history!
Eons: The Longest Intervals of Geologic Time
To represent the rich history of the Precambrian, a division of the geologic time scale longer than the
era, called the eon, was introduced. Four eons, based on the isotopic ages of terrestrial rocks and
meteorites, are now recognized.

HADEAN EON The earliest eon, whose name comes from Hades (the Greek word for "hell"), began
with the formation of Earth 4.56 billion years ago and ended about 3.9 billion years ago. During its
first 650 million years, Earth was bombarded by chunks of material from the early solar system.
Although very few rock formations survived this violent period, individual zircon grains with ages as
great as 4.4 billion years have been found, indicating that Earth had a felsic crust within 200 million
years of its formation. There is also evidence that some liquid water existed on Earth's surface at
about this time, suggesting that the planet cooled rapidly. In Chapter 9, we will explore this early
phase of Earth's history in more detail.

ARCHEAN EON The name of the next eon comes from archaios (the Greek word for "ancient"). Rocks
of Archean age range from 3.9 billion to 2.5 billion years old. The geo-dynamo and the climate
systems were established during the Archean eon, and felsic crust accumulated to form the first
stable continental masses, as we will see in Chapter 10. The processes of plate tectonics were
probably operating by the end of the Archean, although perhaps substantially differently from the
way they did later in Earth's history. Life, in the form of primitive single-celled microorgan-isms,
became established, as indicated by the fossils found in sedimentary rocks of this age.

PROTEROZOIC EON The last part of the Precambrian is the Proterozoic eon (from the Greek words
proteros and zoi, weaning "earlier life"), which spans the time interval from 2.5 billion to 542 million
years ago. By the beginning of this eon, the plate tectonic and climate systems were working much
as they do today. Throughout the Proterozoic, organ-isms that produced oxygen as a waste product
(as plants do today) increased the amount of oxygen in Earth's atmosphere. We will explore the
early evolution of life and its effects on the Earth system in Chapter 11.

PHANEROZOIC EON The start of the Phanerozoic eon is marked by the first appearance of shelly
fossils at the beginning of the Cambrian period, now dated at 542 mil-lion years ago. The name of
this eon—from the Greek ykaneros and zoi ("visible life")—certainly fits, because it comprises all
three eras originally recognized in the fossil record: the Paleozoic (542 million to 251 million years
ago), the Mesozoic (251 million to 65 million years ago), and the Cenozoic (65 million years ago to
the present).

Perspectives on Geologic Time

In the dusty sheep country of far western Australia stands a small promontory of ancient red rocks
called the Jack Hills (Figure 8.16). Geologists have pulverized truckloads of these rocks to isolate a
few sand-sized crystals of zircon. By measuring the lead-206 and lead-207 isotopes generated by the
radioactive decay of uranium-238 and uranium-235, as described above, they have identified one
small crystal fragment with an age of 4.4 billion years—the oldest mineral grain yet discovered in
Earth's crust. How can we relate to such a mind-boggling span of time?
Imagine compressing the 4.56 billion years of Earth his-tory into a single year, starting with the
formation of Earth on January 1 and ending at midnight on December 31. Within the first week,
Earth was organized into core, man-tle, and crust. The oldest zircon grain from the Jack Hills
crystallized on January 13. The first primitive organisms appeared in mid-March. By mid-June, stable
continents had developed, and throughout the summer and early fall, the biological activity of
evolving life increased the concentration of oxygen in the atmosphere. On November 18, at the
beginning of the Cambrian period, complex organisms, including those with shells, appeared. On
December 11, reptiles evolved, and late on Christmas Day, the dinosaurs became extinct. Modern
humans, Homo sapiens, did not appear on the scene until 11:42 P.M. on New Year's Eve, and the
most recent ice age did not end until 11:58 P.M. Three and a half seconds before midnight,
Columbus landed on a West Indian island, and a couple of tenths of a second ago, you were born!

Recent Advances in Timing the Earth System

We have seen that the time scales of geologic processes are not uniform, but vary from seconds to
billions of years. We must therefore use a variety of methods for timing the Earth system: some to
determine the ages of very old rocks, others to measure rapid changes. New methods for
determining the relative and absolute ages of Earth materials have steadily improved our
understanding of how the Earth system works. To conclude our story of the geologic time scale, we
will describe a few of these recent advances.

Sequence Stratigraphy
Until a few decades ago, geologists had to rely on rocks exposed at outcrops, in mines, and by
drilling to map stratigraphic successions. As mentioned in Chapter 1 (and further described in
Chapter 14), technological innovations in the field of seismic imaging now allow us to see below
Earth's surface without actually digging holes.
 From recordings of seismic waves generated by controlled explosions or by natural earthquakes, we
can construct three-dimensional images of deeply buried structures (Figure 8.17). Seismic imaging of
sedimentary rocks allows geologists to identify sedimentary sequences and map their distribution in
three dimensions, a type of geologic mapping called sequence stratigraphy.

Sedimentary sequences commonly form on the edges of continents; here, sediment deposition by
rivers is modified by fluctuations in sea level. In the example shown in Figure 8.17, sediments were
laid down in a delta where a river entered the ocean. As the sediments accumulated, the delta
advanced seaward. When sea level fell because of continental glaciation, the deltaic deposits were
exposed and eroded. Once the glaciers melted and sea level rose, the shoreline shifted inland, and a
new deltaic sequence began to cover the old one, creating an unconformity.

Over millions of years, cycles such as this one can be repeated many times, producing a complex set
of sedimentary sequences. Because sea level fluctuations are global, geologists can match
sedimentary sequences of the same

Recent Advances in Timing the Earth System 213

over large areas. The relative ages of these sequences then be used to reconstruct the geologic
history of a n, including any regional tectonic uplift or subsidence contributed to sea level changes.
Sequence stratigraphy-has been especially effective in finding deeply buried and gas on continental
margins, such as the Gulf of co and the Atlantic margin of North America.

Chemical Stratigraphy Many sedimentary beds contain minerals and chemi-that identify
them as distinct units. For example, the unt of iron or manganese in carbonate sediments vary from
bed to bed if the composition of seawater changed during precipitation of the carbonate minerals. n
the sediments are buried and transformed into sedimentary rocks, these chemical variations may be
pre-, "fingerprinting" the formations. These chemical fingerprints may extend regionally or even
globally, allowing us to match sedimentary rocks by chemical stratigraphy where no other features,
such as fossils, are available.

Paleomagnetic Stratigraphy
Another technique for fingerprinting rock formations s paleomagnetic stratigraphy. As we saw in
Chapter 1, Earth's magnetic field reverses itself at irregular inter-1s. These magnetic reversals are
recorded by thermonanent magnetization in volcanic rocks, which can be Lazed by isotopic methods.
The resulting chronology of magnetic reversals—the magnetic time scale—allows vs to "replay the
magnetic tape" of seafloor spreading and determine the rates of plate movements, as we saw
Chapter 2. Even more detailed patterns of magnetic reversals can be observed in sediment cores,
and these magnetic fingerprints can be dated using faunal successions. Paleomagnetic stratigraphy
has recently become one of the main methods for measuring sedimentation rates along the
continental margins and in the deep sea. We will discuss paleomagnetic stratigraphy in more detail
in Chapter 14.

Clocking the Climate System

The Pliocene and Pleistocene epochs were times of rapid and dramatic global climate change. We
can chart these climate changes from the isotopes contained in shelly fossils buried in deep-sea
sediments. Deep-sea drilling vessels such as the JOIDES Resolution (see Figure 2.13) have taken cores
from sedimentary beds around the world's oceans. Geologists can use the carbon-14 dating method
to estimate when the shells recovered from these sediment cores were formed, and they can
measure the stable isotopes of oxygen to estimate temperature of the seawater in which the shell-
producing organisms lived.
 The carefully tabulation of both temperature and age estimates for many sedimentary layers has
provided us with a precise record of global climate during the last 5 million years (Figure 8.18). The
record shows a general cooling trend beginning about 3.5 million years ago and the subsequent
development of rapid climate cycles that be-came especially large during the Pleistocene epoch. The
low temperatures during these cycles, which were as much as 8° C below the average present-day
temperature of Earth's surface, correspond to the Pleistocene "ice ages," when glaciers covered
large areas of North America, Europe and Asia.

Repeated cycles of glaciation have occurred with dominant periods ranging from 40,000 to 100,000
years. Shorter-term cycles lasting a few thousand years or less are also evident. The effects of these
climate cycles, such as rises and drops in sea level, can have profound effects on Earth's surface. We
will explore glacial cycles and their causes in more detail in Chapters 15 and 21.

SUMMARY
How do we know whether one rock is older than another?
We can determine the relative ages of rocks by studying the stratigraphy, fossils, and cross-cutting
relationships of rock formations observed at outcrops. According to Steno's principles, an
undeformed sequence of sedimentary beds will be horizontal, with each bed younger than the beds
beneath it and older than the beds above it. In addition, the fossils found in each bed reflect the
organisms that were living when that bed was deposited. Knowing the faunal succession makes it
easier to spot unconformities, which indicate time gaps in the stratigraphic record where no rock
was deposited or where existing rock was eroded away before the next strata were laid down.

How was a global geologic time scale created?

By using faunal successions to match rocks in outcrops around the world, geologists compiled
composite stratigraphic successions, from which they developed a relative time scale. The use of
isotopic dating allowed them to assign absolute ages to the eons, eras, periods, and epochs that
constitute the geologic time scale. Isotopic dating is based on the decay of radioactive isotopes, in
which unstable parent atoms are transformed into stable daughter atoms at a constant rate. By
measuring the amounts of parent and daughter atoms in a sample, geologists can calculate the
absolute ages of rocks. The isotopic clock starts ticking when radioactive isotopes are locked into
minerals as igneous rocks crystallize or metamorphic rocks recrystallize

What are the principal divisions of the geologic time scale?

The geologic time scale is divided into four eons: the Hadean (4.56 billion to 3.9 billion years ago),
Archean (3.9 billion to 2.5 billion years ago), Proterozoic (2.5 billion to 542 million years ago), and
Phanerozoic (542 million years ago to the present). The Phanerozoic eon is divided into three eras,
the Paleozoic, Mesozoic, and Cenozoic, each of which is divided into shorter periods. The boundaries
of the eras and periods are marked by abrupt changes in the fossil record; many correspond to mass
extinctions.
What other methods are now being used to date the geologic record? The
cyclical rise and fall of sea level produces complex sedimentary sequences on continental margins
around the world that can be mapped using seismic imaging techniques and dated using fossils.
Chemical fingerprints and magnetic reversals provide additional information about the ages of
sedimentary sequences. Glacial cycles recorded in sediments can be dated using ice cores taken
from the Antarctic and Greenland ice caps.