1 Stable isotopes, seafloor sediments, and climate

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climatic optimum

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Figure 6.1 Record of 18 O enhancements from fossils of benthic foraminifera, ocean creatures in seafloor sediments, ranging in age from

65 million years ago to the present. The horizontal axis is time before present; the corresponding geological epochs of the Cenozoic era

(Chapter 8) are marked on the figure. The measure of isotopic enrichment, δ 18 O, is by convention the difference of 18 O/16 O in the sample to that of

a known standard, divided by that reference standard 18 O/16 O value. As is common to avoid having to use decimals and zeros, the values are all

multiplied by 1,000. Larger amounts of δ 18 O correspond to larger global ice volume and hence colder conditions, and a rough temperature

calibration is shown on the right-hand axis down to the freezing point of water. Major climate events over this time period are labeled, as are the

onsets and durations of the large Antarctic and northern hemisphere ice sheets. The geologic epochs along the bottom are explained in Chapter 8.

Adapted from P¨ like and Hilgen, 2008.

a



oceans to form clouds. During cold periods on earth – ice ages –

the polar caps were built up and spread outward in the form of

ice sheets. The sources of the ice sheets are storm systems that

dump large amounts of moisture on high-latitude continents in

the form of snow. Therefore, water actually was lost from the

oceans during ice ages and stored as ice in great continental

ice sheets. Because it is the lightest isotope H2 16 O that is preferentially evaporated from oceans, during ice ages the oceans

should be enriched in 18 O. The same kind of enrichment occurs

for 17 O, but its mass difference from 16 O is half that of 18 O, and

so, climate effects are smaller. For that reason, we focus only on

the 18 O/16 O record because it is read more easily.

Again, a storage medium for the oxygen is required and, as

before, tiny sea creatures play the role. To make the calcite shells

or plates such as coccoliths requires oxygen as well as carbon

(CaCO3 ). As with carbon, the oxygen is taken up without regard

for the isotope number, 18 O and 16 O equally. The oxygen source

is water in the ocean. Hence these microscopic shells record the

ratio of 18 O to 16 O in the ocean at the time of their formation.

The measurement for oxygen (and carbon, for that matter) is

available from organisms at both the ocean’s surface and its

depths, because oxygen extracted by deep-sea shellmakers will

record the local 18 O/16 O ratio in the ocean water.

As the organisms die, the calcite shells are deposited on the

ocean floor, buried by progressive sedimentation and become

part of the rock record. Ancient ocean sediments exposed by

geologic events allow the calcitic oxygen to be extracted, the

18

O/16 O ratio measured, and the global temperature determined.

To put these data into a climate chronological sequence, the sediments or surrounding rock layers then must be dated. Figure 6.1

shows an example of ocean temperatures derived from the

oxygen isotopic data for the time period from 65 million years

ago, which covers the demise of the dinosaurs through the golden

age of mammals (see Chapter 19). It should be remembered

that the oxygen isotopic ratio is only an imperfect measure of

ocean temperature, since other effects such as ocean salinity can



alter isotopic ratios. Hence cross-comparison with other types

of data, such as distribution of planet species in the fossil record,

is essential.



6.1.3 Hydrogen

About 10 parts per million (ppm) of water on Earth contains

heavy hydrogen, or deuterium, primarily paired with a light

hydrogen to make HDO (as opposed to normal water, H2 O). As

with the heavier isotopes of oxygen, deuterated water tends to

be preferentially left behind during evaporation of ocean water

near the equator. Thus air masses moving away from the equator

and hence toward colder latitudes are slightly enriched in normal water; that is, they possess somewhat less than the 10 ppm

of HDO that is typical for the ocean. As rain and snow form in

the air mass, the deuterated water is preferentially and progressively removed from the storm system in the precipitation, so that

storms near the poles drop snow that is significantly depleted

in deuterium. The depletion is exaggerated during colder climate episodes relative to warmer, because the drop in temperature from equator to pole is larger during colder times. (This

is checked by mapping the distribution of plant species during

warm times versus ice ages.) Also, the tendency of deuterium to

fall out in the rain and snow is exaggerated at lower temperatures

(Figure 6.2).

The resulting deuterium fractionation has been used to study

the most recent epochs of glacial climate and warm episodes

in between, by sampling the ice sheets that cover Antarctica,

Greenland, and other very cold places. The record in such ice

sheets extends back less than 300,000 years in Earth history, but

it is much more detailed than that in the more ancient seafloor

sediment record of carbon and oxygen isotopes. Colder times are

characterized by more exaggerated deuterium fractionation and

hence greater deuterium depletion in the ice laid down at that

time. Warmer episodes show less deuterium depletion. Because

the ice also contains the oxygen isotopes discussed above, the



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(a) Warmer global temperatures



Decreasing local t

empe

ratur

e



s



“light water” (H2O)



small

temperature

contrast



G

” (HDO)

wa ter

avy

h

“He



heavier HDO

evaporates last,

precipitates first



fractionation

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eva



poration

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ua

to

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North Pole



(b) Cooler global temperatures



Decreasing loc



al tem

p



eratu

r



es



“light water” (H2O)



larger

temperature

contrast



G

O)

ter” (HD

y wa

ea v

h

“H



heavier HDO

evaporates last,

precipitates first



eva



poration



Eq

ua



to

r



fractionation

effects

exaggerated



precipitation

North Pole



Figure 6.2 Progressive depletion of deuterium-bearing water from equatorial ocean to the polar ice sheets, during (a) warmer and (b) colder times.

The steeper equator-to-pole temperature drop under cold climate conditions exaggerates the fractionation relative to that in warmer times. HDO is

shown in black, H2 O in gray.



deuterium and 18 O depletions can be compared to help build confidence in the paleo-temperatures (ancient temperatures) derived

from the core. Chapter 21 discusses the application of stable isotopes to understanding the nature of interglacial warm periods

such as that in which we now live.



are composed of long chains of atoms that preserve well in sea

sediments. The relative abundances of some of these molecules,

called alkenones, reflect sea surface temperature at the time of

their formation. The advantage of this so-called UK 37 index

is that it is not influenced by salinity to the same extent as the

oxygen isotopes.



6.1.4 Other systems

There are many other methods for determining past temperatures on Earth; here we have described only a few widely

applied techniques. Nitrogen isotopic ratios are a good indicator

of phytoplankton productivity, and as well sulfur isotopes provide a range of information on the state of past environments

on Earth. Also, certain organic molecules derived from algae



6.2 A possible temperature history of Earth

from cherts

Oxygen isotopic exchange potentially provides a temperature indicator back through 80% of Earth’s history. Cherts



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Largest values at a given time define the upper

envelope, which may provide (for certain

conditions) the surface temperature



0



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1,000



1,500



2,000



2,500



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Time before present (millions of years)



Figure 6.3 A possible temperature history of the Earth is reflected in the oxygen isotopic ratio in samples of chert. Plotted against time in Earth’s

history is the so-called δ 18 O, which is the difference of the ratio 18 O/16 O measured in the sample to that in a reference standard, normalized by the

reference standard. As is common to avoid having to use decimals and zeros, the values are all multiplied by 1,000. As discussed in the text, an

estimate for the surface temperature of the Earth is given by the upper envelope of the curves. Geologic era, which we introduce in Chapter 8, are

labeled and their boundaries indicated by dashed lines. From Knauth (2005).



are hard rocks, composed largely of very-fine-grained silica.

Silica is formed from silicon and oxygen: SiO2 . It can occur as

bands in limestones, as nodules, or in other physical forms.

Cherts form in a wide range of environments, precipitating

directly out of rivers or ocean waters, or forming from rocks

that are subjected to mild increases in temperature and pressure. Biogenic chert, that is, chert made by organisms such as

sponges or radiolaria that secrete silica, is probably the most

abundant.

Of interest to us here is that the oxygen isotopic content

of the chert bears a definite relationship to that of the environment in which it is made. If precipitation occurs in an

ocean environment, then the 18 O content of the silica decreases

with increasing temperature in a manner that can be quantified in the laboratory. Essentially, the chert, which preserves

very well as a sediment through time, acts to record the ambient water temperature through the oxygen isotopic enhancement during its formation. High temperatures in the water

from which the chert is precipitated lead to lower 18 O/16 O

ratios in the chert, while low temperatures lead to high 18 O/16 O

ratios.

Unfortunately, using cherts as indicators of the surface temperature of Earth is extremely complicated because cherts form

in so many different environments and the 18 O/16 O values may

be altered in ways that have nothing to do with the surface

temperature. In the 1970s geochemists Paul Knauth at Arizona

State University and Donald Lowe at Louisiana State University

attempted to use cherts to determine ancient ocean temperatures

in spite of these difficulties. They argued that, for most (but not

all) types of chert, processes during or after formation would

tend to lower the 18 O/16 O enhancement in cherts relative to the



value obtained during precipitation from ocean waters. Therefore, for a collection of cherts of a given age, the cherts with

the highest 18 O/16 O values should most nearly reflect equilibration with ocean waters during formation. Hence, the cherts with

the highest 18 O/16 O value at a given time provide a measure of

Earth’s ocean temperature.

A relative temperature history of the Earth from cherts is

shown in Figure 6.3 from Knauth (2005). If one calibrates the

oxygen isotopic data with the recent temperatures of 10 to 15 ◦ C,

the drop in 18 O/16 O going back in time implies a temperature of

55 to 85 ◦ C in the first third of the Earth’s history. Particularly

striking is the sharp change in global temperature at about 2.5

billion years ago. This sharp drop in temperature occurs roughly

in the time range where other types of geologic evidence suggest that the Earth experienced at least two episodes of dramatic

global cooling in which ice covered much of the Earth (“snowball Earth”; see Chapter 19). Also, as Knauth (2005) points out,

the earliest life forms, based on study of the relationships in

the genome of organisms existing today (Chapter 12), preferred

environments as warm as those implied by the earliest chert

data.

Some qualifications must be applied to this analysis. First, the

chert samples were formed over a range of latitudes, leading to

the concern that one is mixing latitudinal and time variations

in temperature. As we discuss in Chapter 19, however, warmer

ice-free climates, which have dominated over most of Earth’s

history, experienced much less variation of temperature with latitude than we experience today. The second issue is more serious, and has to do with whether the oceanic value of 18 O/16 O has

really been constant over time. One source of variation are the

episodes of massive glaciation interspersed throughout Earth’s



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history that are mentioned above. Formation of glaciers alters

the baseline 18 O/16 O value in the oceans. Further, the baseline

18

O/16 O abundance may have been lower than today’s during

the first quarter of Earth’s history, based on the chemistry of the

most ancient cherts, which would explain the discontinuity in

18

O/16 O at 2.5 billion years ago without invoking a sharp drop in

temperature. Finally, one must remember that the surface temperature is interpreted to be provided by the chert samples with

the highest 18 O/16 O out of a large range. There is no guarantee

that the uppermost value actually corresponds to the ambient surface temperature, even though this is a reasonable assumption

for the later data. If areas of high geothermal activity were more

prevalent in the ancient past than today, it is more likely that the

older samples, even those with the largest 18 O/16 O, are affected

by, or indeed predominantly reflect the environment of, these hot

spots. That many of the samples reflect conditions in and near hot

spots, such as hydrothermal vents on the ocean floor, is supported

by measurement of silicon isotopic ratios, which vary in the

chert samples over time in a way that suggests alteration in hot

vents.

The chert story illustrates how important it is to be cautious

and skeptical when extracting conclusions from a single type of

data. Unfortunately, other evidence for Earth’s temperatures in

the first quarter of its history is extremely sketchy. The intriguing conclusion from the chert data, that at least portions of

the Earth’s oceans were as warm if not warmer than today, is



59



consistent, at least, with the extensive geologic evidence that

liquid water was stable on Earth at least 3.8 billion years ago.

Although this may not seem surprising, it is of some significance

because stellar models argue that the Sun was much less luminous at that time than it is today. The geologic record tells us

that this early warm period was punctuated (or even terminated)

by glacial epochs, in which ice covered much or perhaps all of

the Earth, beginning about 2.9 billion years ago. Evidently the

Earth’s climate underwent an adjustment that we do not understand, or periods dominated by ice existed earlier (but were not

recorded in the chert data); however, the geologic evidence is

too scant to record this.

We discuss the faint early Sun problem in Chapter 14. For

now, it suffices to note that the interpretation of the oxygen

isotopes in chert as indicating high surface temperatures on the

early Earth create a paradox because they require early Earth to

have been significantly warmer than at present when, in fact, the

Sun was significantly dimmer. Spacecraft images of Mars and

direct measurement of rocks at the Martian surface also indicate

that our neighboring planet was hotter earlier in its history than

it is today.

This dual-planet dilemma regarding dramatic climate change

early in the history of the planets, in the face of the faintness

of the Sun at the time, represents a major puzzle that we must

tackle later in the book.



Summary

Stable isotopes of major elements, that is, isotopes that do not

decay measurably over Earth’s history, can be used to track the

climate history of our planet. In order to use stable isotopes,

the given element must be commonly present in sediments or

life forms, it must have more than one stable isotope whose

separation depends on temperature, the altered isotopic ratio

must be preserved in a time-ordered or datable way for a long

time, and the isotopic ratios must be measurable. For recent climate, isotopes of carbon, oxygen, and hydrogen are available.

Carbon has two stable isotopes, of mass 12 and 13, respectively, and the lighter isotope is preferentially incorporated

from atmospheric carbon dioxide into carbohydrates produced

in plants by photosynthesis. Thus, during warm periods, when

less land is covered by ice and more rainfall occurs, allowing

more plant activity, the ratio of the heavier to the lighter isotope, 13 C to 12 C, is enriched in the atmospheric carbon dioxide.

This record is preserved by shell-forming organisms that take

up the atmospheric carbon for their shells, and upon dying

become part of the sediments on the ocean floor. Oxygen and

hydrogen isotopes record climate based on the difference in



propensity for evaporation between the lighter and heavier

isotopes. Because the temperature drops more steeply from

equator to pole during cold periods relative to warm ones,

preferentially more of the lighter isotope is extracted from the

ocean water in cold times and sequestered at the poles as ice.

Thus, in colder times the enrichment of the heavier isotope in

the ocean water is larger than in warmer times. For oxygen

this is recorded in shells; for hydrogen the record is in the

cores of ice deposited during colder climates and preserved in

Antarctica and elsewhere. An oxygen isotopic record of ancient

climate exists in silicon–oxygen rocks called cherts. The cherts

form by precipitation from ocean water, and the isotopic ratio

of oxygen is altered as a function of temperature during the

precipitation into the mineral phase. The cherts suggest that

temperatures in the ocean were higher in the early history of

the Earth than is the case today; however, the chert record may

be contaminated by a number of factors other than the mean

ocean temperature, including the effect of high-temperature

“hydrothermal” vents.



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Questions

1. Why is it important to use more than one isotopic system to



determine the history of Earth’s surface temperature?

2. What is it about the possible differences between 12 C and

13

C that would lead to a preferential uptake by planets of

the former compared to the latter? Is this a phenomenon of

chemistry? If so, could abiotic chemical processes in, for

example, the atmospheres of Mars (Chapter 15) or Saturn’s

moon Titan (Chapter 16) exhibit the same sort of fractionation of the isotopes? Or is this a possible way to detect

biological versus purely abiotic chemical processes?

3. The carbon and oxygen isotopic ratios in shell-forming

organisms is not entirely independent of oceanic conditions;

rather, these ratios might be altered by the amount of hydrogen carbonate ions (ions with the formula HCO− historically

3



called bicarbonates) in the oceans (see Chapter 14 for a discussion of the chemistry). As the carbonate ions (which have

the formula CO− ) become more abundant in the ocean, the

3

values of 13 C/12 C and 18 O/16 O decrease in the shells. Since

hydrogen carbonate ions are primarily produced from erosion of rock by rainfall, and then ends up in the oceans by

river runoff, what might you predict would be the direction

of this effect given that rainfall is decreased during colder

epochs? How might you correct for this effect in determining

past ocean temperatures from the shells of organisms?

4. In using cherts to determine global temperatures in the past,

how would you test the claim that the oceanic 18 O value has

been constant over Earth’s history?



General reading

Considine, D. M. (ed.) 1983. Cherts. In Van Nostrand’s Scientific

Encyclopedia. Van Nostrand Reinhold, New York, p. 624.

Kasting, J. F. and Kirschvink, J. 2012. Evolution of a habitable

planet. In Frontiers of Astrobiology ed. C. Impey, J. Lunine

and J. Funes. Cambridge University Press, in press.



References

Jouzel, J. and Merlivat, L. 1984. Deuterium and oxygen 18 in precipitation: modeling of the isotopic effects during snow formation.

Journal of Geophysical Research 89, 11,749–57.

Knauth L. P. 2005. Temperature and salinity history of the Precambrian ocean: implications for the course of microbial evolution.

Palaeogeography, Palaeoclimatology, Palaeoecology 219, 53–

69.

Knauth, L. P. and Lowe, D. R. 1978. Oxygen isotope geochemistry of cherts from the Onverwachte group (3.4 billion years),

Transvaal, South Africa, with implications for secular variations in the isotopic composition of cherts. Earth and Planetary

Science Letters 41, 209–22.

P¨ like, H. and Hilgen, F. 2008. Rock clock synchronization. Nature

a

Geoscience 1, 282.



Prahl, F. G. and Wakeham, S. G. 1987. Calibration of longchain alkenones as indicators of paleoceaongraphic conditions.

Nature 330, 367–9.

Shackleton, N. J. 1986. Paleogene stable isotope events. Paleogeography, Paleoclimatology, Paleoecology 57, 91–102.

Spero, H. J., Bijma, J., Lea, D. W., and Bemis, B. E. 1997. Effect of

seawater carbonate concentration on foraminiferal carbon and

oxygen isotopes. Nature 390, 497–500.

Van den Boorn, S. H. J. M., van Bergen, M. J., Nijman, W., and

Vroon, P. Z. 2007. Dual role of seawater and hydrothermal

fluids in Early Archean chert formation: evidence from silicon

isotopes. Geology 35, 939–42.

Vostok Project Members. 1995. International effort helps decipher

mysteries of paleoclimate from Antarctic ice cores. EOS 76,

169.



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7

Relative age dating of cosmic and

terrestrial events: the cratering record

Introduction

Instead, scientists use relative dating techniques to infer time

histories of the moons and planets in the solar system, and

they rely primarily on the record of bombardment, or cratering, of the surfaces of these bodies. We describe this

technique and the physics of cratering in the present chapter. In addition to providing a foundation for inferring key

aspects of the solar system’s history, this discussion provides

a good foundation for the presentation in Chapter 8 of relative age dating on Earth, which relies on geologic processes

other than cratering but for which the principles are much the

same.



The absolute dating techniques of Chapter 5 rely on very precise laboratory analyses of rock samples. For Earth, an abundance of accessible samples exists. However, with respect to

the rest of the solar system, only meteorites, small bits of asteroidal and cometary debris – interplanetary dust particles (IDP),

and samples from the Moon have been delivered to terrestrial laboratories for age analyses. One class of meteorites,

the Shergottites–Nakhlites–Chassigny (SNC), may have been

ejected from Mars by collision with one or several asteroids.

Aside from these cases, we have no known samples of material from large bodies in the solar system and thus cannot date

major geologic events on the surfaces of the bodies in an absolute fashion.



7.1 Process of impact cratering

Impact cratering is a process in which a high-speed projectile

collides with a solid surface, forming an excavated region called

a crater. Impact craters, and the closely related form of craters

caused by massive explosions, such as nuclear detonations, can

be distinguished from those produced by other processes, such

as volcanism or collapse due to groundwater withdrawal, by

their distinctive appearance (Figure 7.1).

Projectiles impact with velocities imparted by virtue of their

orbital motion and the gravitational pull of the target planet.

Impact speeds vary depending on the target planet’s distance

from the Sun and the strength of its gravitational field. Typical

impact speeds onto the Moon, due to the free-fall velocity of

projectiles at 150 million kilometers from the Sun, are 40 kilometers per second, or just short of 100,000 miles per hour.

An automobile hitting a surface at 160,000 km (approx.

100,000 miles) per hour delivers an impact energy a million times higher than if it were in a head-on collision with

another vehicle at 80 km (approx. 50 miles) per hour, that is,

160 km (approx. 100 miles) per hour relative velocity, because

impact energy scales as the square of the velocity. However, it

also scales with the mass, and the bigger craters on planetary



surfaces are formed by impactors that are kilometers in size. The

energy released by just one such impactor, kilometers across, is

equivalent to the release of the world’s entire nuclear arsenal –

at the peak prior to current disarmament – many dozens of times

over! Such an enormous release of energy on a habitable planet

has the capability to transform oceans and atmospheres, and to

destroy life on a planetary scale.

At impact with the ground, the projectile plows into the surface, its energy of motion rapidly converted into heat, and the

impact itself sends shock waves, familiar examples of which are

thunder and sonic booms, into the ground. The ground itself

is compressed and shattered by the enormous temperatures and

pressures of the shock wave. The projectile also is shocked and

shattered. The shock waves travel outward and downward in

the ground in a hemispherical pattern. Nearest the impact, rock

is vaporized or melted; farther away it is pulverized. As the

shock waves travel away from the impact, the ground begins to

rebound toward the center of the hemispherical cavity or crater,

forming a central peak in the case of moderate-sized to large

impact craters. The central peak can form only because the rock

is in a temporary state of being partly molten and partly solid,



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(a)



(c)



(b)



(d)



Figure 7.1 Examples of craters formed by different processes: (a) Caldera at the summit of the Martian volcano Aescraeus Mons (mosaic of

NASA Viking images generated by J. Zimbelman at the Lunar and Planetary Institute); (b) sinkhole near Montevallo, Alabama, 120 meters across,

formed by the action of groundwater (US Geological Survey photo); (c) explosion crater about 250 meters across in Nevada, generated by a

30-kiloton nuclear warhead detonated underground (US Department of Energy); (d) meteor crater, Arizona, a small (1 km diameter) impact crater.



the solid part being so weak that the shock waves moving back

and forth can readily push the material. As the shock waves

dissipate, the central peak remains intact.

The shock waves also raise a rim around the crater as well

as eject material off the sides, this material (ejecta) shooting

into the air as hot molten (liquid) rock, traveling many times

the size of the crater away from the center, forming lines of

smaller secondary craters as well as streaks or rays of material

as it strikes the ground. The impactor itself is obliterated and

becomes a small part of the ejecta.

Figure 7.2 shows the stages of crater formation and the final

shapes of typical small and large craters. There are many variations: small craters do not have well-developed central peaks.

Extremely large impactors send shock waves through deeper

parts of the target’s interior, where the warmer rock or ice can

flow more easily, creating large-scale wave patterns that are

preserved as multiring basins. Mare Oriental on the surface

of Earth’s Moon, Valhalla on Jupiter’s moon Callisto, and Gilgamesh on Jupiter’s moon Ganymede are examples. Craters also

may take on different forms depending on whether ground ice is



present, and the strength of the planetary crust: weak crusts will

cause crater topography to disappear over time, leaving only

ghostly outlines. Finally, erosion by water and subduction of

crust (Chapter 9) have removed most of the craters on Earth,

and left many others barely discernible (Figures 7.3a–f).



7.2 Using craters to date planetary surfaces

Craters can be used to determine how old one surface is relative

to another because the rate of impacts over time is thought to

have declined slowly over the past three-quarters of solar system

history, having decreased quickly prior to that from a much

larger initial rate. Surfaces that are young, that is, which have

been renewed through lava flows, mountain building, erosion

by water, and other geologic processes, will show fewer craters

than surfaces that are much less active, or older. Because of this

we can use the abundance of craters on various surfaces of a

planetary body to determine, in a relative sense, when certain

kinds of geologic processes occurred relative to others. The



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(a)



(b)



excavation and

beginning of

uplift



63



excavation and

beginning of

uplift



central uplift

forms



central uplift and

rim collapse



final crater



peak ring crater



Figure 7.2 Stages in the formation of (a) small and (b) large craters. Modified from Melosh (1989, p. 142) by permission of Oxford University

Press.



freshness of craters, that is, how bright their debris or ejecta

blankets appear and how sharp their features are, provides an

additional refinement to the relative dating process. It is also

possible to compare relative ages from one planet or moon to

another, provided one can calibrate the rate of impacts in one

part of the solar system relative to another.

An example from the Moon provides a classic illustration. The

bright areas of the surface of the Moon are very heavily cratered

regions called the lunar highlands, as revealed by telescopes and

images from lunar orbiting spacecraft (Figure 7.4a). In these

regions the density of craters is so great that craters overlap

with and are superimposed on each other; down to the limit of

resolution on the image, one sees a scene filled with craters.

The dark portions of the Moon, on the other hand, consist

of areas that are smooth and relatively devoid of craters. These

mare (Latin for seas, the seventeenth and eighteenth century

interpretion from telescopic views) show ample evidence of

craters that have been partly covered or obliterated by the material that makes up the smooth, dark surfaces (Figure 7.4b). The

simplest interpretation is that the mare are lowland basins that

were flooded by lavas sometime after an early, heavy bombardment of the Moon occurred. The flooding obliterated most of the

craters, leaving a fresh surface on which some remains of old

craters can be seen, and a few small, new craters were formed

by impacts after the lava solidified.

Explorations by the Apollo astronauts from 1969 to 1972

returned nearly 400 kg (900 lbs) of moon rocks from mare

and highland regions. Radioisotopic techniques, described in

Chapter 5, were used to provide absolute dates for the solidification of these rocks from original molten materials. The lunar

highlands are old, with rocks dating as old as 4.5 billion years.

The mare deposits typically are 4.2 billion to 4.3 billion years

old, significantly younger than the highlands.

The age estimates based on the cratering density on the lunar

surface are confirmed by the absolute dating of mare and highlands provided by rock samples. It then would appear possible

to use crater densities on other worlds, not accessible for sample

collection at present, to construct chronologies as well. The most



straightforward chronology involves determining relative ages

of events on a surface, that is, which event preceded another.

This simply requires counting craters as well as looking for evidence of craters partly obliterated by geologic processes. More

difficult is to try to assign actual dates, which requires assuming

that the lunar crater density and ages based on Moon rocks can

be transferred directly to other bodies in the solar system.



7.2.1 Relative ages of events on a planetary surface

Impact craters can be used to determine the relative ages of geologic features on a planetary surface. Two examples of this are

shown, one from Mars and one from Jupiter’s moon Ganymede,

using images from Viking and Galileo missions in Figures 7.5a

and 7.5b, respectively. In the case of Mars, the geologic features

of interest are channels that clearly were cut by water, but today

are dry along with the rest of the planet. Are the features young

or ancient? Was the climate wet up through recent times, such

that life might have evolved to an advanced stage?

Examination of the Viking images such as that in figure 7.5a

reveals that the Martian channels typically are overlain by impact

craters, some fairly substantial in size. Other regions of the Martian surface have far fewer craters, and hence we can say that

the channels are, relatively speaking, ancient. Determining a

more exact age requires tying the cratering rate to some absolute timescale. At the same time, we know that the channels are

not among the oldest features, either, because many cut through

craters that must therefore be older than the channels. A chronology can be assembled in which channel formation occurs after

formation of the oldest Martian terrains but before a number of

other geologic events that are recorded in the surface.

Ganymede, the third of Jupiter’s four giant moons (Io is closest

to Jupiter, and then Europa, Ganymede, and Callisto), shows

lines in its spectrum typical of water ice. However, the mass of

the planet is too heavy given its volume (mass over volume is

density) to be pure water ice. The best guess, based on models

of solar system formation, is that the heavier component is a

silicate, or common rocky material containing silicon, oxygen,



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(a)



(b)



(c)



(e)



(d)



(f)



Figure 7.3 Varieties of impact craters: (a) large multiring basin, Mare Oriental, on the Moon; (b) classic large crater, Copernicus, with central

peak, on the Moon; (c) smaller lunar craters without peaks; (d) pedestal crater on Mars, formed by melting of ground ice during impact. (e) relaxed

craters, or palimpsests on Ganymede (Voyager image); (f) eroded crater on Earth, now comprising Lake Manicouagan, Quebec, Canada (see color

version in plates section). Photos (a) through (f) are courtesy of NASA.



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RELATIVE AGE DATING OF COSMIC AND TERRESTRIAL EVENTS



65



(a)



(b)



Figure 7.4 Two very different terrains on the Moon: (a) the lunar highlands show craters of all sizes filling all available surface space; (b) the lunar

mare regions are smooth, dark plains with a few fresh craters and remains of large craters, in various states of preservation, which were present

when lavas flooded the lunar surface.



(a)



(b)



Figure 7.5 (a) NASA/Viking image of Martian surface cut by channels; (b) NASA/Galileo image of dark and light terrains on Jupiter’s moon,

Ganymede.



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THE MEASURABLE PLANET



magnesium, some iron, and other elements. Therefore, unlike

our own Moon, Earth, Venus, Mercury, and Mars, which are

made up mostly of silicon-bearing rock and metal, Ganymede

is half-rock, half-ice. This is true for Callisto, as well, but not

Europa and Io: they are both mostly rock, although Europa has

an outer veneer of ice and possibly liquid water.

One might expect, from terrestrial experience, that the ice

might behave differently in an impact than rock. In fact, the

pressures and temperatures in the hypervelocity impacts we have

been describing are so large that there is little difference. Furthermore, temperatures in the distant outer solar system, where

Jupiter and its moons reside, are so low that ice behaves much

like rock as a material making up the solid crusts, or outer layers, of Ganymede and Callisto. The surface temperature near

the equator of these moons is typically 165 K, very far below

the ice melting point of 273 K.

Images of Ganymede reveal two types of surfaces: a dark,

heavily cratered terrain, and a bright, lightly cratered terrain.

The paucity of craters on the latter surface immediately suggests

that it is a younger feature, perhaps ice that has been extruded

from the interior along cracks and flowed outward. In places, it

is possible to see where a crater on the older dark terrain has

been partially obliterated by the new material. It is also possible

to tell something about how the cracks and new material formed

by looking at distortions in partially preserved craters along the

edges of the bright terrain. The difference in brightness between

the dark and light terrains remains a matter of speculation; silicates and perhaps some carbon-bearing materials are well mixed

with the water ice, perhaps dating back to the original formation

of Ganymede.

The ability to learn something about the sequence of events

on a surface by looking at crater densities is a tool of primary

importance in solar system studies. It is a new development

of a much older technique applied to Earth geology to look at

superposition of layers to assemble a history of a given region.

On Earth, water and geologic activity have effectively erased

the cratering record, so that the use of craters as a geologic tool

was a novel idea that did not come into its own until planetary

exploration began some three decades ago.



7.2.2 Absolute chronology of solar system events

Relative age dating is limited in the amount of information

derived. Ideally, one wants to assign ages to events on the surfaces of planets and moons so as to understand their history and

ultimately that of the solar system. Imagine how limited our

own understanding of the history of human cultures would be

if we only knew the order of events, but not their antiquity or

duration.

In the case of Earth’s geologic history, even before radioisotopic dating provided reliable dates, estimates of ages could be

made on the basis of notions of the accumulation rate of sediments, debris brought from high to low places by the action of

water. Early work tended to overestimate the rates of sedimentation and hence produced a compressed timescale relative to what

is accepted today based on radioisotopic determinations. With

the help of radioisotopic dating, the rates of geologic processes



are now better understood and calibrated, such that indirect

dating techniques such as sedimentation are enhanced as tools

in assembling the history of Earth.

The situation for an absolute chronology from planetary cratering is similar in that radioisotopic dating has been used to

construct a chronology for Earth’s Moon, which then has been

applied, with caveats, to other solar system bodies. The Moon is

the only body for which radioisotopic dating of terrains of varying crater density can be performed; Earth’s cratering record

is too sparse. (It is not possible to determine unequivocally

from what part of the Martian surface the SNC meteorites were

derived; hence they cannot help calibrate the cratering record on

Mars.)

The oldest parts of the Moon, the highlands, have by far the

largest number of craters; the younger mare possess the least.

This is consistent with the decreasing population over time of

debris in orbits around the Sun. Theories of planet formation,

which we discuss in Chapter 10, hold that the planets were

assembled from smaller pieces of rock and ice through relatively

low-speed collisions that allowed the pieces to stick together. In

the final phases of this process, most of this protoplanetary

material was perturbed by close encounters with the planets

into highly elliptical orbits, guaranteeing that any subsequent

collisions with the planets would be at high speeds, producing

craters. Over time this remnant debris of planet formation was

swept up by the planets, so that the available impactor population

has decreased dramatically from the beginning to the present

day.

A simple law governing the rate of impacts over time, consistent with the sweep-up picture described above, and with

the lunar cratering record, has the inverse exponential form

shown in Figure 7.6. The curve is characterized by a very steep

decrease initially, as large amounts of material are swept up

by the nearly fully grown planets, followed by a transition to

a slowly decreasing rate of impacts. The cratering record on

the Moon tells us when the transition occurs between these two

regimes. Further, it provides information about the tail-off in

impacts at later times, though with limited capability because

of the paucity of new craters. More difficult to discern is the

precise steepness of the early curve, because the cratering rate

was so high that lunar highland surfaces are completely covered with craters: new impacts simply obliterate all or part of

old ones and only a lower limit on the ancient cratering rate is

accessible.

The dating of Moon rocks fixes the transition in the cratering

curve at roughly 3.8 billion to 4.0 billion years before present;

the period of intense cratering before that is called the Late

Heavy Bombardment, referring to the tail end of the planetformation (accretion) process. Interestingly, the oldest whole

rock samples on Earth date back to roughly the same time. We

know that this does not represent the age of Earth because the

rocks are rather evolved, showing the action of liquid water

on their chemistry and texture; additionally, meteorites record

much earlier dates back to 4.56 billion years before present.

Instead, Earth was simply too active geologically at earlier times

to preserve older rocks and, as we see in later chapters, had

little or no continental land mass on which such rocks could be

preserved.



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