Boggs S. Principles of Sedimentology and Stratigraphy

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15.4 Chronocorrelation

537

Note that chronocorrelation on the basis of physical event stratigraphy re

quires that event beds can be recognized and aced laterally in outcrop or that

they can be correlated on the basis of distinctive lithology (lithocorrelation). Be

cause they were produced as a result of an event that took place rapidly, lithocor

relation also results in chronocorrelation. Of course, the actual ages of event beds

must be established by the radiometric dating techniques discussed in the preced

ing section.

Biologic events include episodes of punctuated evolution, mass extctions,

mass mortalities (caused, for example, by major ash fall to a basin), and rapid

immigration and emigration (Kauffman, 1988; Wa llister, 1996). Some of the tech

niques and problems of ronocorrelation by biologic events are discussed in the

preceding chapter on biostratigraphy.

Event Correlation Based on T nsgressive-Regressive Events

A different approach to event correlaon is represented by local correlation based

on position within a transgressive-regressive succession or cycle (Ager, 1993b).

According to Ager, event correlation in this case is based on e correlation of cor

responding peaks of smetric sedimentary cycles that are presumed to be syn

chronous. The events represented in this type of correlation are the result of

transgressions and regressions that may represent either worldwide, simultane

ous, eustatic changes in sea level or more local changes owing to upft, subsi

dence, or fluctuation in sediment supply.

The principle of correlation based on transgressive-regressive events is illus

trated in Figure 15.6. The deposits formed during any transgressive-regressive

cycle contain one particular time plane that represents the time of maximum in

undation by the sea, that is, the time at which water depth was greatest at any par

ticular locality. Rocks lying stratigraphically below this time plane were deposited

during transgression and those above during regression. This time plane can be

identified by use of fossil data to determine depth zonation and maximum water

depth at various localities, as illustrated  Figure 15.6. The position of e time

plane can be established also from liologic evidence by determining in the verti

cal stratigraphic section at each locality the position within the section where the

rocks are symmetrically distributed with respect to the most basinward facies pre

sent. A surface connecting the most basinward rocks in each of the vertical sec

tions defines the approximate position of the time plane and thus the

time-stratigraphic correlation between the sections. Figure 15.7 further illustrates

the method. Note from this illustration how time-equivalent points on the cycle

Landward

Localities

c D E

Seaward

Figure 15.6

Time correlation by position in a transgressive

regressive cycle. The line connecting points of

deepest-water condition is a time line. [Aer

lsraelski, M. C., 1949, Oscillation chart: Am . Assoc.

Petroleum Geologists Bull., v. 33, Fig. 3, p. 98.]

538

Chapter 15 1 Chronostratigraphy and Geologic me

Figure 15.7





w E

. :

.-.

·:

.·

Transgressive-regressive cycle sedimentation and event correlation in the

Eocene of the Isle of Wight in southern England. [From Ager, D. V., 1993, The

nature of the stratigraphical record, 3rd ed., Fig. 7.2, p. 100. Reprinted by

permission of john Wiley & Sons Ltd.

 CROSS - BEDDED SANDS

§ MINATED BEDS

� GUCONITIC CYS

are related, resulting in a correlation in which glaucotic clays at the east end of the

succession are equated to laminated beds at the west end. Correlation is expressed,

as Ager (1993b, p. 101) puts it, in terms of degrees of "marineness." Correlaon in

this manner can be considered to be a part of sequence stratigraphy (Chapter 13).

Correlation by Stable Isotope Events

Variations in the relative abundance of certain stable, nonradioactive isotopes in

marine sediments and fossils, referred to as stable isotope geochemistry (e.g., Val

ley and Cole, 2001), can be used as a tool for chronostragraphic correlation of

marine sediments. Geochemical evidence shows that the isotopic composition of

oxygen, carbon, sulfur, and strontium in the ocean has undergone large fluctua

tions, or "excursions," in the geologic past-fluctuations that have been recorded

in marine sediments. Because the mixing time in the oceans is about 1000 years or

less, marine isotopic excursions are considered to be essentially isochronous

throughout the world. Variations in isotopic composition of sediments or fossils

allow geochemists to construct isotopic composition cues that can be used as

stratigraphic markers for correlation purposes. To be useful for correlation, fluctu

ations in isotopic composition must be recognizable on a global scale and must be

of sufficiently short duration to show up as a shift on isotopic composition curves.

Also, stratigraphers must be able to fix the relative stratigraphic position of these

fluctuations in relation to biostratigraphic, paleomagnetic, or radiometric scales.

Of the various potentially useful isotopes, oxygen isotopes seem most nearly to

meet these requirements and have proven to be parcularly useful for chronos

tratigraphic correlaon of Quaternary and late Tertiary sediments. Carbon, sulfur,

and strontium isotopes are also useful for correlating rocks of certain ages.

Oxygen Isotopes

The natural isotopes of oxygen are listed in Table 15.3. Most of the oxygen in the

oceans occurs as oxygen-16. Oxygen-18 is much rarer (about 0.2 percent of total

Atomic

Element

number

s 16

Source: CRC Handbook of Chemistry and Physics.

Isotope

15.4 Chronocorrelation

539

Relative Isotopic

abundances (%)

99.789

0.037

0.204

98.89

1.11

95.0

0.76

4.22

0.014

0.56

9.86

7.02

82 .56

oxygen), but it is present in measurable amounts. The ratio of

180

160

in the

ocean at any given time in the past is built into contemporaneous marine carbon

ate minerals and the calcium carbonate shells of marine organisms as a permanent

record of the isotopic composition of the ocean at those times. Fluctuations in oxy

gen isotope ratios in the ocean with time thus show up in the geologic record as

fluctuations in e isotopic ratios of these marine carbonates and fossils. Classifi

cation of deep-sea sedents on the basis of oxygen isotope ratios in the shells of

calcareous marine orgisms, particularly foraminifers, has given rise to a new

stratigraphy for Quaternary sediments. This stratigraphic method is commonly

referred to as oxygen isotope stratigraphy. It was first used by Emiliani (1955),

who studied the isotopic composition of foraminifers in deep-sea cores and used

oxygen isotope ratios to subdivide the core sediments. Oxygen isotope stratigra

phy has now developed into a major tool for correlating Quaternary and late Te r

tiary marine successions, as explained below.

The 180/160 ratio in biogenic marine carbonates reects both the tempera

ture and the 180

160 ratio of the water in which these carbonates formed. The re

lationships

of ocean paleotemperature (T) to oxygen isotopic composition has

been shown by Shackleton (1967) to be

T(0C)

16.9 - 4.38(de dw) + 0.10(de dw

(15.1)

where de the equilibri oxygen isotope composition of calcite and dw = oxygen

isotope

composition of the water from which the calcite was precipitated. The de

and dw notations refer not to the actual oxygen isotopic abundances in calcite and

water but to the per mil (parts per thousand) deviation of the 180/160 ratio in cal

cite and water from that of an arbitrary standard. A coonly used standard for

oxygen isotopes in the past was the University of Chicago PDB standard, where

PDB refers to a particular fossil belemnite from the Pee Dee Formation of South

Carolina. More commonly now, the isotope composition of ocean water (Standard

540

Chapter 15 I Chronostratigraphy and Geologic me

Mean Ocean Water, or SMOW) is used as a stan

ard (e.g., Coplen, Kendall, and

Hopple, 1983). The per mil deviation from the standard, referred to as a

o, is ex

pressed by the relationship

[(

sample - e

0) standard]

o =

x 1000 (15.2)

standard

Oxygen isotope stratigraphy is based on the fact that a

o values in biogenic ma

rine carbonates reflect both e temperature and the isotopic composition of the

water from which e calcite precipitates. These factors are both, in tum, a func

tion of the climate. When water evaporates at the surface of e ocean, e lighter

0 isotopes

are preferentially removed in the water vapor, leaving the heavier

o 

the ocean. This isotopic fractionation process thus causes water vapor to be

depleted of

O with respect to e seawater from which it evaporates. When

water vapor condenses to form rain or snow, the water containing heav� oxygen

will tend to precipitate first, leaving the remaining vapor depleted in

o com

pared to the initial vapor. Thus, the precipitation that falls near the coast and runs

back quickly to the ocean will contain heavier oxygen than that which falls in e

interior of continents or in polar regions, where it returns more slowly to the

ocean. The

0 ratio of precipitates also correlates with the air temperature.

The colder e air, the liter the rain or snow (Odin, Renard, and Grazzini, 1982).

For example, the overall average oxygen isotope composition of seawater is

.28 %o (per mil); however, the precipitation that falls in e crests of the Green

land Ice Sheet is about -35%o and in relavely accessible parts of the Antarctic

Ice Sheet it is as negative as -58 %o .

The

0-depleted moisture that falls in polar regions is locked up as ice on

land and is thus prevented from quickly retuing to e ocean. Because of this re

tention of light-oxgen water in the ice caps, the ocean becomes progressively en

riched in

o as

O-depleted ice caps build up during a glacial stage. Marine

carbonates that precipitate in the ocean during a glacial stage, particularly bio

genic carbonates such as foraminifers, will be enriched in

0 relative to those that

precipitate during times when the climate is warmer and ice caps are absent, or

are

much smaller, on land. Changes in e a

o content of biogenic marine calcite

us reflect changes in the volumes of ice on land and concomitant changes in sea

leveL That is, sea level drops as ice masses build up on land during glacial e�isodes

and rises when continental ice masses melt durg interglacial stages. The a

0 val

ues of seawater track these changes, becoming higher (more posive values) dur

ing glacial stages when heavy oxygen is concenated in the ocean and lower (more

negative) during terglacial stages as melng of polar ice caps retus light-oxy

gen water to the oceans. ese principles are illustrated in Figure 15.8.

Decrease  temperature of the seawater in which biogenic calcite precipi

tates also causes an increase  e a

o values that are built to the calcite. Thus,

during glacial periods both decrease in temperature of ocean water and changes

in isotopic composition of ocean water owing to buildup of ice caps on the conti

nents combine to crease the a

o content of biogenic calcites. Conversely, melt

ing of polar ice caps, with consequent return of light-oxygen water to the oceans,

and increase in ocean temperature will be reflected in a decrease in 8

o values in

marine biogenic carbonates.

Dierent kinds of marine organisms tend to incorporate somewhat different

ratios of oxygen isotopes into their shells (fractionate oxygen isotopes to dierent

degrees), as indicated in Figure 15.9. Therefore, to evaluate changes in oxygen iso

topes in the ocean as a function of time requires that we analyze the same kind of

fossil organism in rocks of different ages. Planktonic foraminifers are the most

common fossil used in oxygen isotope studies of this kind.

Q (%o}

�

0.5 -1 .0 -1 .5





Hermatypic

corals

 Full Interglacial: sea level high stand



Te rmination of glaciation: rapid isotopic

and sea level change



Full Glacial: ice volume maximum, sea level miminum,

0 enrichment

in oceans

Period of slowly increasing ice volume and

sea level lowering

Last Interglacial: arrows denote

sea level changes

_

Te rmination of glaciation: rapid isotopic and

sea level change

Full Glacial: -150,000 years before present

Green algae

Freshwater

limestone

Shallow-water

marine

limestone

\ Deep-sea

limestone

Shallow-water

molluscs and

foraminifers

15.4 Chronocorrelation 541

Figure 15.8

Schematic illustration of the relationship between

continental glaciation and the o

80 content of

ocean water during the last 150,000 years. Note

that

0 values become less negative (more

heavy oxygen) during stages of continental glacia

tion and lowered sea level. [After Williams, D. F.,

I. Lerche, and  Fult 1988, Isotope chronos

tratigraphy-Theory and methods: Academic

Press, Fig. 30, p. 58. Reproduced by permission.]

Figure 15.9

Distribution of o180 and o1

C values in various

types of marine carbonates. (After Milliman, J. D.,

1974, Marine carbonates. Fig. 19, p. 33. Reprint

ed by permission of Springer-Verlag.]

542

Chapter 15 I Chronostratigraphy and Geologic Time

Figure 15.10

Oxygen-isotope stratigraphy

of Pleistocene cores from

Deep Sea Drilling Project

(DSDP) sites in the Pacific and

Atlantic oceans. Note the

numbered isotope stages,

which

can be correlated from

one core to another. M/B

refers to the Matuyama/Brun-

hes polarity chrons. [After

Wei, W., 1993, calibration of

upper Pliocene-lower Pleis-

tocene nannofossil events

with oxygen isotope stratigra-

phy: Paleoceanography, v. 8.,

Oxygen isotopes in sediment cores of late Tertiary to Quateary age show

numerous 8180 maxima and minima. Figure 15.10 shows an example of oxygen

isotope values in three Pleistocene cores om widespread localities in e Pacific

and Atlantic oceans (Wei, 1993). The numbers on the isotope curves are oxygen

isotope stage numbers (e.g., Kennett, 1982; Ruddiman et al., 1989; Raymo et al.,

1989). Where available, the magnetostratigraphic chrons are shown also. Once e

isotope stages in a core have been identified and numbered, they can be correlated







:



40 







Site 647

5 4 3

2 0

Site 593

2 0

Fig. 4, p. 91, published by the

American Geophysical Union.]

120

160E

160W

120

15.4 Chronocorrelation

543

to e same isotope stages in other cores across the world ocean. These isotopic

events appear to be related to e Milankovich orbital-climate cycles discussed 

Chapter 12. us, they provide a record of fourth- and fifth-order cycles that are

presumably driven by changes in climate related to changes in Earth's orbital para

meters. These orbital chges cause uctuations in the intensity of solar radiation

reaching Earth at different latitudes, which, in tum, result in alteate accumula

tion and melting of continental ice sheets, producing rise and fall of sea level.

Carbon Isoto

Carbon-12 and carbon-13 are e nonradioactive isotopes of carbon. Carbon-12 is

much more abundant than carbon-13 and makes up most of the carbon  seawater

(Table 15.3). The isotopic

C ratio can  expressed  terms of per l deviaon

( 8

C) from e PDB standard, just as oxygen isotope ratios are expressed. The 8

values  marine carbonates reflect e

C rao of C0

dissolved in deep ocean

water;

is ratio, in tum, reects the source of carbon  C0

• Carbon dioxide dis

solves in the ocean by interchange with the atmosphere, and it is generated also by e

decay of organic matter that originates both in the ocean and on ld. Organisms pref

erenally incorporate light carbon (

C); therefore, carbon dioxide derived om de

caying orgac matter is shaly depleted of

C compared to at derived om e

atmosphere. us, water runo om the continents (where soil organic matter is

abundant) brings orgc-ri waters with low

C ratios into the ocean, sii

cantly lowering the 8

C content of surface ocean waters near the continents. (Note

from Fig. 15.9 the low 8

C values in freshwater carbonates deposited in lakes.)

Another factor that influces e 8

C contt of ocean water, and us the 8

content in the shells of marine organisms that live in these waters, is the sidence

me of deep-water masses in the oce. Carbon-13 is depleted in deep-water masses

that have long residence times near the ocean bottom, owing to oxidaon of low-8

marine orgc matter that sinks om the surface. Oxidation of this low-a

C organic

matter leads to producon of low-a

C dissolved bicarbonate (HC0

-). Respiration

by bottom-dwelling organisms also aepartly causes a decrease in a

C of deep bot

tom waters (Kennett, 1982). If low-a

C bottom water is later circulated to the surface

in some maer, carbonate-secreng organisms will build this low-a

C isotope ratio

into their shells.

The a

c content of ocean water is also related to the primary productivity of

photosynthesizing marine organisms (e.g., diatoms). During times of high pro

ductivity (large numbers of organisms present), the rate of removal of light carbon

(

by these organisms is high. Selective removal of the light carbon

causes surface ocean waters to be relatively enriched in

e, resulting in positive

values of a

C. The light carbon is delivered to the seafloor when organisms die,

whe it is reoxidized (Holser and Magaritz, 1992). During times of low primary

productivity, this end is reversed and surface waters tend to have negative a

values. It has been suggested, for example, that the dramatic decrease in a

c val

ues across the Cretaceous-Tertiary boundary (see Fig. 15.11) is the result of marked

decrease in marine photosynthesizing organisms owing to a catastrophic mete

orite impact (e.g., Hsu et al., 1982).

Because the a

c in the calcareous shells of marine organisms is a function of

the a

c content of the waters in which they live, changes in the o

c content of

fossil marine organisms indicate changes in ocean water masses. Abrupt decreas

es in the a

c in fossil marine calcareous organisms may reflect changes in prima

ry marine productivity, as suggested, or changes in deep ocean paleocirculation

d upwelling patterns that caused low-a

c deep waters to spread upward and

outward into other parts of the ocean. Or such decreases may reflect changes in

544 Chapter 15 I Chronostratigraphy and Geologic Time

:



l





C (%o PDB)

Site 528

Site 525A

Site 527

Site 529

C (%o PDB)

Figure 15.1 1

Carbon isotope stratigraphy of bulk carbonate sediments from Deep Sea Drilling Project

(DSDP) Sites 525A, 527, 528, and 529 in the Atlantic Ocean about 800 km off the coast of

Africa. The sites are about 40 to 70 km apart. [After Shackleton, N. ]., and M. A. Hall, 1984,

Carbon isotope data from Leg 74 sediments, in Moore, T. C., Jr., and P. D. Rabinowitz

et al., Initial Reports DSDP 74, Washington, U.S. Gov. Printing Office., Fig. 1, p. 614.]

surface circulation patterns that brought low-8

C surface ocean waters from con

tinental margins into deeper basins. Significant increases in the total biomass pro

duced on the continents during any particular geologic time interval could cause

an increase in runoff of low-8

C to the oceans, an increase that may be reflected

also as an episode of low-8

C surface water. Increased rates of erosion of organic

rich sediments, such as dark shales and limestones, on land could produce much

the same effect of increasing runoff of low-8

C in organic matter to the ocean.

These abrupt changes in circulation patterns or organic carbon runoff from the

continents may have affected the area of a single ocean basin, such as the Pacific,

or in some cases the entire ocean. On the other hand, increased rates of sediment

burial in the ocean may have the opposite effect of removing fine organic matter

containing low-8

C from interaction with seawater. This removal would have the

effect of increasing o

C in ocean water. Several increases in 8

C that took place 

the middle Miocene between about 12-18 Ma may be associated with exceptionally

rapid burial of organic carbon and lowered levels of atmospheric C02 (e.g.,

Woodruff and Savin, 1991).

Inasmuch as changes in the 8

C content of the ocean are reflected in the 8

content of marine calcareous organisms, these isotopic events or "excursions" can

be used for correlation, regardless of their exact causes. For example, several

major changes in the Tertiary 8

C record include (1) a pronounced negative 8

event across the Cretaceous-Te rtiary boundary, (2) a positive event from the early

15.4 Chronocorrelation

545

Paleocene into the late Paleocene, and (3) a major a

C decrease across the Pale

ocene-Eocene boundary (Fig. 15.11). There is also a gradual but significant a

deterioration from the mid-Miocene to the Pleistocene (e.g., Williams, Lerche, and

Full, 1988, p. 54). Many smaller-scale fluctuations in the Tertiary a

C record are

also present. These carbon isotopic excursions are essentially synchronous events

that can be correlated over wide areas of the ocean in DSDP and ODP cores.

Major carbon isotope events are also present in the older sedimentary record.

For example, large changes in carbon isotope content of carbonates have been

measured at the Precambrian/Cambrian boundary and the Permian/Triassic

boundary (Magaritz, 1991).

Sulfu r Isoto

Sulfur has four stable isotopes (Table 15.3); sulfur-32 is the most abundant, fol

lowed by sulfur-. The

S ratio is used in most stratigraphic studies involv

ing sulfur isotopes and is expressed  terms of a

S, which is per mil deviation of

the

S ratio relative to a meteorite standard, troilite (a FeS mineral) from the

Canyon Diablo meteorite. Figure 15.12 shows the a

S values in various materials

relative to the Canyon Diablo standard (CDT).

The major means of sulfur isotope fractionation in the oceans is bacterial re

duction of sulfate (S0

-) in seawater to sulfides (H

S, Hs- , HS0

-). Bacterial re

duction of dissolved seawater sulfate at the sediment-seawater interface causes

isotopic fractionation of the sulfate, thus enriching the remaining seawater sulfate

in a

S by about +20%o and depletion in the reduced sulfide by about -9%o (e.g.,

Schopf, 1980; Canfield, 2001). Precipitation of evaporites from dissolved marine

sulfates introduces an additional fractionation (�+1.65%o), causing the a

S of

evaporites to be higher than that of dissolved marine sulfates. Other minor factors

that can influence the a

S content of seawater include oxidation of bacterial

which produces sulfates depleted of

S relative to original sulfates, and local em

anations of sulfate or sulfide through volcanic activity.

Marine sulfates in the present ocean have a mean a

S of about +21 %a,

however, the a

S of ancient marine evaporites ranges from about + 10 to +30

(Fig. 15.13). On the basis of sulfur isotope ratios in ancient evaporite deposits, it

appears that the sulfur isotope ratios in the surface waters of the world ocean have

undergone major changes, or excursions, at various times. These major excursions

are characterized by sharp rises in a

S in the surface waters of the world ocean fol

lowed by significant drops. The early Paleozoic exhibits high a

S levels, dicative

Canyon Diablo meteorite

(standard)

Marine sulfates, recent

Marine sulfates, ancient

�Seawater sulfate

 Rainwater sulfate

Sedimentary sulfides

�

Figure 15.12

S values of marine sulfates shown rela

tive to that of the Canyon Diablo mete

orite. Values for seawater sulfate, rainwater

sulfate, and sedimentary sulfides are shown

for comparison. [Data from Degens, 1965.]

546

Chapter 15 I Chronostratigraphy and Geologic Time



100

200



6 300



�

400

Figure 15.13

Sulfur isotope age curve for Phanerozoic marine evaporites. Dashed

portions of the curve indicate lack of data. The dashed vertical line

indicates the o

S composition of marine evaporites in the modern

ocean. [After Holser, W. T., M. Magaritz, and j. Wright, 1986, Chem

ical and isotopic variations in the world ocean during Phanerozoic

time, in Walliser, 0. (ed.), Global bio-events, Lecture Notes in Earth

Sciences, v. 8, Fig. 4, p. 69, Springer-Verlag.]

500

600



.

' 

-



-



+10

+20

48, 0/00 COT

+30

of a high net flux of sulfide from ocean to sediments; 8

5 drops to a minimum in

the Permian, and it returns to intermediate levels in the Mesozoic (Holser, Maga

ritz, and Wright, 1986). Many of these sulfur isotope excursions appear to have af

fected the ocean worldwide.

Chemical events characterized by sharply increased 8

5 may be caused by

catastrophic mixing of deep

5-rich brines with surface waters. Brines generated

by evaporite deposition are stored in deep basins. Underneath the brines, bacteri

al reduction of sulfates to form pyrite builds up a store of brine heavy in sul

fate. Catastrophic mixing of these

-rich brines with surface waters, owing to

destruction of the storage basin by tectonism, causes a sharp rise in the 8

5 of sur

face ocean waters and consequently in the evaporite deposits formed from these

surface waters (e.g., Holser, 1977; 1984). Gradual decrease in the 8

5 of surface

ocean waters with time after a catastrophic event is attributed to on-land erosion

predominantly sulfide materials into the ocean in an amount that exceeds evap

orite deposition (e.g., Claypool et al., 1980).

In any case, these sulfur isotope excursions con.titute a sulfur isotope age

cur

ve because each catastrophic chemical event occurred within a very short in

terval of geologic time. Each major event thus represents a synchronous stra

graphic marker that can be correlated in marine evaporite deposits from one area

to another. Some of the events can be correlated on a global basis. Thus, they pro

vide an important method for inteational chronostratigraphic correlation of

evaporite deposits, which commonly cannot be correlated by other means because

they do not contain fossils or other datable materials. On the other hand, only a

few well-defined, correlatable points are present on the sulfur-isotope curve; thus,

sulfur isotopes are less useful overall for correlation purposes than are oxygen and

carbon isotopes.